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The geobiology of fluvial, lacustrine, and marginal marine carbonate microbialites (Pleistocene, Miocene, and Late Triassic) and their environmental significance
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The geobiology of fluvial, lacustrine, and marginal marine carbonate microbialites (Pleistocene, Miocene, and Late Triassic) and their environmental significance
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by
_____________________________
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THE GEOBIOLOGY OF FLUVIAL, LACUSTRINE, AND MARGINAL MARINE
CARBONATE MICROBIALITES (PLEISTOCENE, MIOCENE, AND LATE TRIASSIC)
AND THEIR ENVIRONMENTAL SIGNIFICANCE
by
Yadira Ibarra
A Dissertation Presented to the
FACULTY OF THE USC GRADUATE SCHOOL
UNIVERSITY OF SOUTHERN CALIFORNIA
In Partial Fulfillment of the
Requirements for the Degree
DOCTOR OF PHILOSOPHY
in
GEOLOGICAL SCIENCES
December 2014
!
! i
Table of Contents
List of Figures…………………………………………………………………………………………………….vi
List of Tables……………………………………………………………………………………………………...ix
Acknowledgements………………………………………………………………………………………………x
Chapter 1: Microbial carbonates as spatiotemporal geobiological records ............................ 1
Introduction ................................................................................................................................. 1
Spring-Associated, Fluvial Microbialites as Environmental Indicators ..................................... 1
Stromatolites as Environmental Indicators During Mass Extinctions ........................................ 3
Lateral Continuity and Scales of Control on Microbialite Morphology ..................................... 4
References ................................................................................................................................... 5
!
Chapter 2: Were fossil spring-associated carbonates from Santa Barbara, CA deposited
under an ambient or thermal regime? ........................................................................................ 7
Abstract ....................................................................................................................................... 7
Introduction ................................................................................................................................. 8
Geological And Environmental Context Of Study Site ............................................................ 10
Methods.................................................................................................................................... 10
Facies descriptions ................................................................................................................ 10
Carbonate isotopic analyses .................................................................................................. 11
Water isotope analyses .......................................................................................................... 11
Results ....................................................................................................................................... 12
Facies descriptions ................................................................................................................ 12
Mesostructure ........................................................................................................................ 13
Microstructure observations of light brown and dark brown bands ..................................... 14
Carbonate isotopic analyses .................................................................................................. 15
Water isotopic analyses ......................................................................................................... 15
Discussion ................................................................................................................................. 15
! ii
Oocardium stratum calcite biosignature ............................................................................... 15
Environmental conditions associated with Oocardium calcite deposition ........................... 16
Oxygen Isotope Thermometry .............................................................................................. 18
Carbon Isotopes .................................................................................................................... 21
Significance for climate studies in southern California ........................................................ 23
Conclusions ............................................................................................................................... 24
References ................................................................................................................................. 25
Figures and Tables .................................................................................................................... 33
!
Chapter 3: Fluvial tufa evidence of Late Pleistocene wet intervals from Santa Barbara,
California, U.S.A ......................................................................................................................... 43
Abstract ..................................................................................................................................... 43
Introduction ............................................................................................................................... 44
Tufa evidence for pluvial ...................................................................................................... 45
Geologic And Environmental Setting ....................................................................................... 46
Spring Carbonate Facies ........................................................................................................... 47
Fluvial carbonates ................................................................................................................. 47
Perched carbonates ................................................................................................................ 48
Methods..................................................................................................................................... 49
Sample collection .................................................................................................................. 49
Age control ............................................................................................................................ 49
Carbonate isotopic analyses .................................................................................................. 52
Results ....................................................................................................................................... 53
Age Control ........................................................................................................................... 53
Carbonate stable isotopes ...................................................................................................... 54
Discussion ................................................................................................................................. 54
Comparison of perched and fluvial deposits ......................................................................... 54
Conclusions ............................................................................................................................... 61
References ................................................................................................................................. 61
Figures and Tables .................................................................................................................... 70
! iii
!
Chapter 4: Microfacies of the Cotham Marble: A tubestone carbonate microbialite from
the Upper Triassic, Southwestern United Kingdom ................................................................ 79
Abstract ..................................................................................................................................... 79
Introduction ............................................................................................................................... 80
Lithology and Environmental Setting ....................................................................................... 81
Methods and Approach ............................................................................................................. 82
Results ....................................................................................................................................... 83
Mesostructure of Vertical Cross Sections ............................................................................. 83
Mesostructure of Horizontal Cross-Sections ........................................................................ 85
Microstructure ....................................................................................................................... 85
Discussion ................................................................................................................................. 89
Morphogenesis: Previous Interpretations ............................................................................. 89
Morphogenesis: New Interpretations .................................................................................... 90
Relevance to the End-Triassic Mass Extinction ................................................................... 99
Conclusions ............................................................................................................................. 101
References ............................................................................................................................... 101
Figures and Tables .................................................................................................................. 111
!
Chapter 5: A widespread microbial carbonate response across the end-Triassic mass
extinction, southwestern United Kingdom ............................................................................. 124
Abstract ................................................................................................................................... 124
Introduction ............................................................................................................................. 125
Stratigraphy and Environmental Setting ................................................................................. 126
Methods................................................................................................................................... 127
Results ..................................................................................................................................... 127
Geochemistry ...................................................................................................................... 127
Petrography ......................................................................................................................... 128
Discussion ............................................................................................................................... 128
Microbial Facies .................................................................................................................. 128
! iv
Carbon isotopes ................................................................................................................... 129
Significance of microbialite and associated biogeochemical observations ........................ 130
Conclusions ............................................................................................................................. 133
Acknowledgements ................................................................................................................. 133
Refernces Cited ....................................................................................................................... 134
Figures and Tables .................................................................................................................. 141
Chapter 6: Lateral continuity of multi-scale stromatolite morphology: Implications for
assessing the dominant scales of control ................................................................................. 149
Abstract ................................................................................................................................... 149
Processess and Scales of Control ............................................................................................ 151
Lateral Continuity ................................................................................................................... 153
Case Studies ............................................................................................................................ 155
Barstow Formation Lacustrine Tufa ................................................................................... 155
Upper Triassic Cotham Marble ........................................................................................... 156
Discussion of Results from this Study .................................................................................... 158
Broader Implications/Significance .......................................................................................... 160
Conclusions ............................................................................................................................. 162
References ............................................................................................................................... 162
Figures and Tables .................................................................................................................. 169
!
Chapter 7: Multiscale controls on the formation of lacustrine tufa from the Middle
Miocene Barstow Formation, California ................................................................................ 175
Abstract ................................................................................................................................... 175
Introduction ............................................................................................................................. 176
Geologic Setting ...................................................................................................................... 178
Previous Work on the Barstow Tufas ..................................................................................... 179
Megascale Distribution of Tufas ......................................................................................... 179
Geochemical Analyses and Dating ..................................................................................... 179
Methods................................................................................................................................... 180
Sample preparation ............................................................................................................. 180
! v
Stable isotopic analyses ...................................................................................................... 180
Results ..................................................................................................................................... 181
Mesostructure ...................................................................................................................... 181
Microstructure ..................................................................................................................... 181
Stable isotope results ........................................................................................................... 183
Discussion ............................................................................................................................... 183
Interpretation of Microfabrics ............................................................................................. 183
Multiscale Controls on Tufa Formation .............................................................................. 189
Implications for Early Earth Microbialites ......................................................................... 191
Conclusions ............................................................................................................................. 192
References ............................................................................................................................... 193
Figures and Tables .................................................................................................................. 203
! vi
List of Figures
Chapter 2
Figure 1. Geologic setting of spring-associated carbonates.. ........................................................ 33!
Figure 2. Facies of the spring carbonate deposits.. ....................................................................... 34!
Figure 3. Mesostructure of the spring carbonate fabrics from the four facies in Fig. 2. .............. 35!
Figure 4. Meso and microstructure of the spring carbonate fabric ............................................... 36!
Figure 5. Microstructure of the light brown bands.. ..................................................................... 37!
Figure 6. Photomicrographs of the Oocardium stratum calcite microstructure.. ......................... 38!
Figure 7. Box and whisker plots of calculated temperature estimates comparing the light brown
and dark brown bands. .......................................................................................................... 39!
Figure 8. Carbonate stable carbon and oxygen isotope results. .................................................... 39!
Figure 9. Stable isotope plots versus distance from the boxed spring. ......................................... 40!
Chapter 3
Figure 1. Study site. ...................................................................................................................... 70!
Figure 2. Schematic of carbonate facies modified from Viles et al., 2007.. ................................. 71!
Figure 3. Multiscale facies of the fluvial carbonates. ................................................................... 72!
Figure 4. Multiscale facies of the perched carbonates. ................................................................. 73!
Figure 5. Representative samples used for dating analyses.. ........................................................ 74!
Figure 6. Cross plot of carbonate carbon and oxygen stable isotope values of the fluvial and
perched deposits. ................................................................................................................... 75!
Figure 7. Regional comparisons. ................................................................................................. 76!
Chapter 4
Figure 1. Map of the southwestern United Kingdom showing the locations of the Cotham Marble
microbialites ........................................................................................................................ 111!
Figure 2. The five most common phases of the Cotham Marble. ............................................... 112!
Figure 3. Scans and photomicrographs of L1. ............................................................................ 113!
Figure 4. Bedding plane view of the Cotham Marble ................................................................. 114!
Figure 5. Scan and photomicrographs of D1 .............................................................................. 115!
Figure 6. Putative microfossils in the dendrolites ....................................................................... 116!
! vii
Figure 7. Energy dispersive spectrometry analyses of filamentous microstructures. ................. 116!
Figure 8. Putative Tasmanites phycomata .................................................................................. 117!
Figure 9. Laminations of L2 ....................................................................................................... 118!
Figure 10. Growth phase D2 ....................................................................................................... 119!
Figure 11. Thin section photomicrograph of a sample from Pinhay Bay ................................... 120!
Figure 12. Photomicrographs showing complex microbial mat branching ................................ 120!
Figure 13. Box and whisker plot of percent total organic carbon ............................................... 121!
Figure 14. Various spar and microspar-filled ovoid structures ................................................... 121!
Chapter 5
Figure 1. Upper Triassic stratigraphy of the Southwestern United Kingdom ........................... 141!
Figure 2. Regional Map ............................................................................................................. 142!
Figure 3. Cotham Marble microbialite facies ............................................................................ 143!
Figure 4. Microfacies of the Upper Cotham Member from Lavernock Point, South Wales. .... 144!
Figure 5. Cross plot of carbonate δ
13
C and δ
18
O of the Cotham Marble. ................................... 145!
Chapter 6
Figure 1. Expected lateral continuity relationship between two stromatolite subunits according
to distance decay ................................................................................................................. 169!
Figure 2. In situ macro and mesostructure of Barstow Formation tufa ..................................... 170!
Figure 3. Mesostructure of several Barstow Formation subunits. ............................................. 171!
Figure 4. Thin section photomicrographs of Barstow Formation subunits. .............................. 172!
Figure 5. Upper Triassic Cotham Marble microbialites ............................................................. 173!
Chapter 7
Figure 1. Geologic map of study site. ........................................................................................ 203!
Figure 2. Outcrop images of tufa mound at the Owl Campground tufa-bearing site.. .............. 204!
Figure 3. Mesostructure of tufa subunits ................................................................................... 205!
Figure 4. Mesoscopic textures ................................................................................................... 206!
Figure 5. Microstructure of the weakly laminated, porous regions ........................................... 207!
Figure 6. Laminoid fenestrae ...................................................................................................... 208!
! viii
Figure 7. Oriented filamentous microfossils .............................................................................. 209!
Figure 8. Dense filamentous micrite fabric from the porous, weakly laminated regions. ......... 210!
Figure 9. Cross sections of round molds .................................................................................... 211!
Figure 10. Microfossils of the porous regions ........................................................................... 212!
Figure 11. Photomicrographs of banded carbonate ................................................................... 213!
Figure 12. Mesostructure of four subunits with the bands aligned. ........................................... 214!
Figure 13. Photomicrographs of three tufa subunits with the bands aligned. ............................ 215!
Figure 14. Carbonate δ
13
C and δ
18
O cross plot. .......................................................................... 216!
Figure 15. Carbonate stable isotopes of carbon and oxygen by fabric type .............................. 217!
! ix
List of Tables
Chapter 2
Table 1. Stable isotopic compositions of δ
18
O and δ
13
C (‰ VPDB) for carbonate from their
corresponding facies ............................................................................................................. 41!
Table 2. Stable isotopic compositions of δ
18
O and δD of modern water from Zaca Spring ........ 42!
Chapter 3
Table 1. IRSL measurement values of carbonate samples from the perched cascade core. ......... 77!
Table 2.
14
C age data for carbonate and organics from the fluvial and perched deposits ............. 77!
Table 3. Stable isotopic compositions of δ
18
O and δ
13
C (‰ VPDB) for carbonates from their
corresponding deposit. .......................................................................................................... 78!
Chapter 4
Table 1. TOC measurement results for the dendrolite and fill fabrics ........................................ 122!
Chapter 5
Table 1. Stable isotopic compositions of δ
18
O and δ
13
C (‰ VPDB) for carbon and oxygen from
their corresponding site and microbialite layer. .................................................................. 146!
Table 2. Stable isotopic composition of δ
13
C
org
(‰) of microbialite samples from their
corresponding site and layer. .............................................................................................. 147!
Chapter 6
Table 1. Potential processes of control on stromatolite morphology at various scales. ............ 174!
Chapter 7
Table 1. Stable isotopic compositions of δ
18
O and δ
13
C (‰ VPDB) for carbonate from their
corresponding fabric type. .................................................................................................. 218!
! x
Acknowledgements
I thank my dissertation committee Frank Corsetti, Sarah Feakins, Dave Bottjer, Dave
Caron, and Will Berelson (honorary member) for invaluable discussions and guidance
throughout the course of my PhD. My experience at USC has been exciting, enriching, and
transformative thanks, above all, to my PhD advisor and mentor, Frank Corsetti. I’ve learned
from Frank by watching him lead. Whether it was in the field, in the classroom, or by standing
over his shoulder as he peered through the microscope. Frank offered me the freedom to let
curiosity guide me while also always being available with keen insight and sound advice. I also
wish to thank Will Berelson for his ever-present support, motivation, and constructive critiques.!
My work on the Zaca carbonates was made possible thanks to help and feedback from
Sarah Feakins, who has also been integral in my academic development providing guidance,
advice, and encouragement. I wish to thank Matt Kirby for help in the field and collaboration on
the Zaca project, Ed Rhodes for IRSL dating assistance, and the Zaca Lake Retreat Staff for field
access. Collaboration, discussions, and enthusiasm from Sarah Greene and Dave Bottjer led to
work presented in Chapters 4 and 5. I also thank Sarah Greene and Mike Lewis for their
hospitality during my trips to Bristol, the folks at the Bristol Museum and Art Gallery for access
to museum specimens, and Ramues Gallois and Mike Lewis for field assistance during my time
in the UK.
I thank Miguel Rincon and Nick Rollins for lab assistance and Alonso Lopez, Mike
Cheetham, Audra Bardsley, Luther Beegle, Kris Zacny, and Gale Paulsen for field assistance at
Zaca − despite the bears and poison oak! I wish to also thank the BSWRC, Arlin, Emily, Ms.
Tambara, former and current members of the USC SACNAS Chapter, and members of the
Bottjer and Corsetti labs past and present for friendship and support. Thanks additionally to the
folks of the 2008 New Mexico Tech REU Program − Tom Kieft, Penny Boston, and Sean
Faulkner who sparked my interest in Geomicrobiology, and Lisa Majkowski and Mike Pullin
who introduced me to SACNAS. The USC Earth Sciences Department, the Geological Society
of America, the American Philosophical Society, NSF, and NASA provided funding and travel
support for this research.
Finally, but most importantly, I thank my family. My parents Maria and Humberto
Ibarra, and my sisters Erica, Nena, Yoli, and Gabby for their love and encouragement. Thanks
Sofia, Becky, and Nicolás, I hope to be able to teach you as much as you’ve all already taught
me.
! 1
Chapter 1:
Microbial carbonates as spatiotemporal geobiological records
INTRODUCTION
Microbialites are organosedimentary deposits that accrete as a result of a benthic
microbial community trapping and binding detrital sediment and/or forming the locus of mineral
precipitation (Burne and Moore, 1987). The formation of microbialites is generally thought to
result from a combination of processes (intrinsic and extrinsic) that act on the accreting substrate
over a range of spatial and temporal scales. However, one of the most challenging tasks in the
study of ancient microbialites is assessing the relative importance of the various processes that
control their formation, complicating our ability to interpret their meaning in the rock record.
Understanding the detailed mechanisms and processes that ultimately govern microbialite
formation is an important area of sedimentary research for (1) their role as macroscopic
manifestations of microbial life making them key targets as biosignatures and (2) their
sedimentary nature make them possible repositories of proxy information about their
depositional environment. In this dissertation, three case studies of microbial carbonates are
presented with examples from terrestrial and marine settings that utilize morphological, textural,
and geochemical observations to address, in each case, their geobiological significance.
Observations are then drawn from each case study to arrive at a potential approach that may
assist in unraveling the dominant processes of control on microbialite morphology.
SPRING-ASSOCIATED, FLUVIAL MICROBIALITES AS ENVIRONMENTAL
INDICATORS
Ambient temperature spring-associated carbonates that form subaerially are referred to as
fluvial tufas (Ford and Pedley, 1996). The CO
2
associated with the groundwater that feeds tufa
! 2
deposits is meteoric in origin (derived from the atmosphere and soil zone), leading some workers
to refer to these deposits as meteogene spring carbonates (Viles and Goudie, 1990). Carbonate
deposition results from infiltration, dissolution, saturation, subsurface transport, emergence, and
precipitation (Carthew et al., 2003). As rainwater percolates through the soil horizon, it becomes
enriched in soil CO
2
. With continued subsurface transport, the CO
2
-rich water reacts with the
carbonate bedrock causing the carbonate to dissolve and enriching the groundwater with calcium
and bicarbonate ions. As groundwater flows through the groundwater aquifer it dissolves the
limestone bedrock and the fluids will become supersaturated with respect to calcium carbonate.
During episodes of accelerated groundwater flow, water will emerge at the surface and descend
into areas of lower elevation. Carbonate formation will result from the release of CO
2
, which
can be caused by variations in water temperature, air temperature, turbulence, discharge size, and
biology.
Fluvial tufa deposits only form when the water table is sufficiently drowned to (1) sustain
surface water flow and (2) promote dissolution in the subsurface to saturate the water in
bicarbonate (Pentecost, 2005). Given that tufa growth is dependent on water availability
(precipitation), which is a direct consequence of wet climatic conditions, the presence of fossil
tufa is often used to infer wetter climate in areas that today are characterized by more arid
conditions (e.g., Cremaschi et al., 2010). Fluvial tufas therefore constitute at least two
significant avenues of study: (1) their rapidly lithifying nature offer the opportunity to explore
them from a geomicrobiological perspective, potentially offering insight into the role of benthic
microbial communities in carbonate precipitation and (2) their dependence on sufficient rainfall
make fluvial tufas possible terrestrial archives of hydrologic balance, and thus important
complementary records to other paleoclimate proxies.
! 3
Previously undescribed spring-associated carbonate deposits located near Zaca Lake in
Santa Barbara County, CA are investigated from a geomicrobiologic and paleoenvioronmental
perspective. In chapter two, the multiscale (micro- to macroscopic) facies distribution of the
carbonates is explored to reveal clues about their nature of deposition (ambient temperature or
cold water), and chapter three addresses the regional paleoenvironmental significance of the
carbonate deposits.
STROMATOLITES AS ENVIRONMENTAL INDICATORS DURING MASS
EXTINCTIONS
Marine microbialites used to be a conspicuous component of subtidal marine
environments in the early Earth, before the advent of animal life. However, in the Phanerozoic
the number of microbialite forms in normal marine settings began to decline (Awramik and
Sprinkle, 1999). Although microbialites occur in some modern marine settings, they are
morphologically distinct from their Proterozoic counterparts. Modern microbialites are coarse-
grained and poorly laminated, whereas Proterozoic microbiliates are primarily composed of fine-
grained laminated and/or thrombolitic textures (Awramik and Riding, 1988; Grotzinger and
Knoll, 1999). Throughout the Phanerozoic, widespread deposits of fine-grained microbialites are
largely a byproduct of environmentally restricted settings (e.g., alkaline lakes, spring-associated).
During times of environmental crisis, however, there is a noted increase in microbialite
abundance in normal marine environments oftentimes occurring in the immediate aftermath of
environmental and ecological disturbance, where microbialites expand to occupy normal marine
subtidal settings previously inhabited normal marine fauna affected by the extinction (Schubert
and Bottjer, 1992; Mata and Bottjer et al., 2011). In such instances, microbialites may serve as
critical geobiologic and geochemical repositories of environmentally perturbed systems.
! 4
The Upper Triassic beds of the Southwest United Kingdom contain laterally extensive
(~2,000 km
2
) carbonate microbialite deposits found stratigraphically at or near the End-Triassic
extinction horizon. Following a traditional ‘actualistic’ interpretation of microbialites as
indicators of harsh environments where metazoan grazers are restricted, previous investigators
have suggested deposition in a restricted lagoon (Radley et al., 2008). However, given that non-
actualistic conditions occur during extinction events, brings into question the microbialites
potential relevance to the extinction. Chapter four highlights new details about the microbialites
that address their morphogenesis, and in chapter five the microbialites are investigated from a
regional and stratigraphic perspective, exploring their potential relevance to the End-Triassic
mass extinction.
LATERAL CONTINUITY AND SCALES OF CONTROL ON MICROBIALITE
MORPHOLOGY
To better understand the utility of microbialites as paleoenvironmental archives of
ancient environments, it is important to address the relative roles of local (biology, mineral
growth/dissolution, etc.) and regional (e.g., climate) processes on their morphology. Recent
work has focused attention on sub decimeter scale morphological features, drawing preferential
focus on local controls on microbialite formation, and potentially masking the role of large-scale
processes. More tools are needed to enhance our understanding of the possible role(s) of
nonlocal, regional controls on microbialite morphogenesis. Chapter six features an approach to
the study of microbial (or abiotic) carbonates that uses lateral continuity together with multiscale
observations of microbialite fabrics—and that incorporates observations from the preceding
chapters—to explore the potential morphological imprint of various scales of control on
microbialite morphogenesis. Finally, in chapter seven lacustrine carbonate microbialites from
! 5
the Middle Miocene Barstow Formation located in the Mojave Desert Southern California are
investigated as a test case using a multiscale, lateral continuity approach to unravel the dominant
scales of control on their formation.
REFERENCES
Awramik, S.M., and Riding, R., 1988, Role of algal eukaryotes in subtidal columnar stromatolite
formation: Proceedings of the National Academy of Sciences of the United Sates of
America, v. 85, p. 1327-1329.
Awramik, S.M., and Sprinkle, J., 1999, Proterozoic stromatolites: The first marine evolutionary
biota: Historical Biology, v. 13, p. 241-253.
Burne, R.V., and Moore, L.S., 1987, Microbialites: Organosedimentary deposits of benthic
microbial commnunities: Palaios, v. 2, p. 241-254.
Carthew, K.D., Drysdale, R.N., and Taylor, M.P., 2003, Tufa deposits and biological activity,
Riversleigh, Northwestern Queensland, In Roach, I.C. (ed.), Advances in Regolith:
Proceedingsof the CRC LEME Regional Regolith symposia. CRC LEME, Bentley,
Australia, p. 55-59.
Cremaschi, M., Zerboni, A., Spötl, C., and Felletti, F., 2010, The calcareous tufa in the Tadrart
Acacus Mt. (SW Fezzan, Libya) An early Holocene palaeoclimate archive in the central
Sahara: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 287, p. 81-94.
Ford, T.D., and Pedley, H.M., 1996, A review of tufa and travertine deposits of the world: Earth
Science Reviews, v. 41, p. 117-175.
Grotzinger, J.P., and Knoll, A.H., 1999, Stromatolites in Precambrian Carbonates: Evolutionary
Mileposts or Environmental Dipsticks?: Annual Review of Earth and Planetary Sciences,
v. 27, p. 313-358.
! 6
Mata, S.A., and Bottjer, D.J., 2011, Microbes and mass extinctions: Paleoenvironmental
distribution of microbialites during times of biotic crisis: Geobiology, v. 10, p. 3-
24,10.1111/j.1472-4669.2011.00305.x.
Pentecost, A., 2005, Travertine: Berlin, Springer, 460 p.
Radley, J.D., Twitchett, R.J., Mander, L., and Cope, J.C.W., 2008, Discussion on palaeoecology
of the Late Triassic extinction event in the SW UK: Journal of the Geological Society,
London, v. 165, p. 988-992.
Schubert, J.K., and Bottjer, D.J., 1992, Early Triassic stromatolites as post-mass extinction
disaster forms: Geology, v. 20, p. 883-886.
Viles, H.A., and Goudie, A.S., 1990, Tufas, travertines and allied carbonate deposits: Progress in
Physical Geography, v. 14, p. 19-41.
! 7
Chapter 2:
Were fossil spring-associated carbonates from Santa Barbara,
CA deposited under an ambient or thermal regime?
ABSTRACT
A previously undescribed succession of currently-inactive spring-associated carbonates
located near Zaca Lake, Southern California, was investigated to determine the nature of
deposition (ambient temperature or hydrothermal water, as both are found within the region).
The carbonate deposits are up to ~1 m thick and formed discontinuously for over 200 m in a
narrow valley between two ridges that drain Miocene Monterey Formation bedrock.
Depositional facies along the presently dry fluvial path include barrage deposits, narrow fluvial
channels, and cascade deposits. The carbonates are mesoscopically banded and contain
ubiquitous micro- to macrophyte calcite encrusted fabrics. All of the depositional facies contain
alternating bands (~.05 mm to 5 mm thick) of dark brown and light brown isopachous calcite;
the dark brown bands are composed of dense isopachous bladed calcite, whereas the light brown
bands are composed of bundles of calcite tubules interpreted as the biosignature of the desmid
microalgae Oocardium stratum. Oxygen isotope thermometry utilizing modern water δ
18
O
values from the piped spring reveal depositional water temperature estimates that collectively
range from ~11−16 ºC. Stable isotope carbon values exhibit a mean δ
13
C value of −9.01 ±
0.62‰ (1σ, n=27). Our petrographic and geochemical data demonstrate that (1) inactive
carbonates were likely sourced from ambient temperature water with a strong soil-zone δ
13
C
signal, (2) the Oocardium calcite biosignature can be used to infer depositional temperature and
flow conditions, and (3) the occurrence of extensive carbonates (especially the presence of a
perched cascade deposit) indicate the carbonates formed when conditions were much wetter.
! 8
INTRODUCTION
Terrestrial carbonates that form via spring activity are sensitive to the environment in
which they form, potentially serving as valuable archives of hydrology, water chemistry, biology
and climate of their local depositional setting (Chafetz et al., 1991; Fouke et al., 2000; Pentecost,
2005; Andrews, 2006; Cremaschi et al., 2010; Sanders et al., 2011). Spring-associated carbonate
deposits are commonly classified based on the temperature and origin of the carrier CO
2
of the
water from which they are sourced (e.g., Ford and Pedley, 1996). The term travertine refers to
carbonate deposits that form from high-temperature spring water, where the CO
2
is sourced from
hydrothermal fluids, whereas the term tufa has been assigned to ambient temperature carbonate
spring precipitates, where the CO
2
is sourced from local soils and the atmosphere (Ford and
Pedley, 1996). Their distinct hydrologic origins thereby provide different information on the
environmental setting of deposits that are no longer active. The presence of travertine, for
example, has implications for nearby hydrothermal/volcanic activity (Hancock et al., 1999;
Fouke et al., 2000), whereas the groundwater-fed nature of tufa deposits results in a strong
influence by local climate (Andrews, 2006).
In order to accurately interpret depositional information from a given fossil/inactive
terrestrial carbonate succession, we must first determine the hydrological regime under which the
carbonates formed (ambient temperature, thermal, or a mix of thermal and ambient water).
Degree of lithification, macrophyte encrustations, and organic carbon content allow travertine to
be distinguished from tufa (Ford and Pedley, 1996; Minissale et al., 2002; Gandin and
Capezzuoli, 2008; Capezzuoli et al., 2014). However, poor textural preservation, similar
sedimentological facies, and interlayering/mixing of tufa and travertine in sites where (1)
hydrothermally sourced waters cool with extensive lateral transport and/or (2) where thermal and
! 9
ambient waters from the same site mix to produce interlayering facies of travertine and tufa (e.g.,
Capezzuoli et al., 2008; Pedley, 2009) may complicate our ability to fully decipher depositional
information from fossil/inactive deposits. For these reasons, geochemical analyses (e.g., stable
isotopes) are often used to infer the hydrologic nature of inactive spring carbonates (Gonfiantini
et al., 1968; Szulc and Cwizewicz, 1989; Guo et al., 1996; Minissale et al., 2002; Andrews and
Brasier, 2005; Kele et al., 2008). Nonetheless, terrestrial carbonates are highly susceptible to
recrystalization, possibly compromising geochemical signatures and thus primary structures are
desirable.
A potential complementary approach to verify the hydrology of deposition is to examine
the micromorphology of the deposit. It is well-established that microorganisms can contribute to
the formation and alteration of terrestrial spring carbonates (Freytet and Verrecchia, 1998;
Riding, 2000; Golubić et al., 2008; Pedley, 2009; Arp et al., 2010; Manzo et al., 2012). Much of
the focus on terrestrial carbonates has been on the interplay between the physicochemical and
biological mechanisms that contribute to carbonate precipitation. Active carbonate-depositing
springs serve as natural laboratories where we can study the extent to which microorganisms
might influence the precipitation of carbonates and other minerals. By examining mineral
phases/structures associated with specific taxa (e.g., Freytet and Plet, 1996; Freytet and
Verrecchia, 1998), we may be able to constrain environmental signals from dormant/inactive
deposits presuming they have not yet become diagenetically altered beyond recognition. Here
we combine a multiscale facies approach (sensu Shapiro, 2000) with analyses of stable isotopes
to determine the depositional regime (ambient or thermal) of previously undescribed
fossil/inactive spring carbonates located in Santa Barbara County, Southern California as both
cold springs and hot springs have been reported in the surrounding area (USGS, 1995).
! 10
GEOLOGICAL AND ENVIRONMENTAL CONTEXT OF STUDY SITE
Spring-derived carbonate deposits are located about 2 km upstream of Zaca Lake in Santa
Barbara County, California, USA (Fig. 1). In 1911 a natural spring was boxed and piped to
provide water for consumption (Fig. 1B) such that carbonate deposition downstream has not
been observed in the modern. Carbonates crop out discontinuously for over 200 m along the
stream grade bed of a narrow valley that is bound by two ridges composed of the carbonate-rich
Miocene Monterey Formation (Hall, 1981). The ridges define the relatively small catchment that
drains into Zaca Lake, one of only a few naturally-occurring lakes in Southern California.
Two distinct units of spring-related carbonate growth have been observed. The first spring
carbonate succession, the focus of the majority of the work described here, was deposited
discontinuously for approximately 200 m along the stream grade bed of the valley (Fig. 1B).
The width of the spring carbonate transect varies from about 1 m to approximately the width of
the valley floor (~15 m) and ranges in thickness from 0 to about 1 m. A second carbonate unit
(approximately 15 m in lateral extent and ~2 m thick) occurs perched upon the slope of the north
ridge about 10 m above the modern stream grade bed (Fig. 1B). The timing of carbonate
deposition is discussed in detail in Chapter 3.
METHODS
Facies descriptions
In our approach, we describe the macro- to micro- characteristics (sensu Shapiro, 2000)
of four distinct facies (Fig. 1B) along the spring carbonate transects. Carbonate rock samples
were obtained from the topmost (~10−20 cm) part of the four sections where there was good
surface exposure and samples were easily accessible. Samples were slabbed, polished, and
scanned for mesostructural studies (cm scale). Microstructural observations were carried out via
! 11
light microscopy of thin sections. Complementary thin-section rock pieces were etched with
diluted HCl and examined further using a scanning electron microscope (SEM). Mineralogy was
determined via X-ray diffraction (XRD) at the Los Angeles Museum of Natural History. Here,
we focus on the microfabric of a particular lamination that is conspicuous in outcrop and hand
sample.
Carbonate isotopic analyses
Isotopic analyses of carbonate oxygen and carbon were conducted on an Elementar
Americas Inc. (Micromass Ltd) Isoprime stable isotope ratio mass spectrometer (IRMS) with a
multi-prep/carbonate device and dual inlet at the Stott Laboratory at the University of Southern
California. Samples were drilled from polished hand sample specimens after careful inspection
for recrystalization via thin section analyses. Samples are measured relative to CO
2
reference
gas calibrated against the NBS-19 (δ
18
O value +2.20‰, δ
13
C value +1.95‰) carbonate standard,
which allows for normalization to the 2-point VPDB-LVSEC isotopic scale. The precision of
this determination is better than 0.06‰ and 0.04‰ (1σ, n = 20) for δ
18
O and δ
13
C, respectively.
A working standard (carbonate, δ
18
O −1.88‰, δ
13
C value +2.07‰) monitors precision during
the course of the run to 0.06‰ and 0.04‰ (1σ, n = 43) for δ
18
O and δ
13
C, respectively.
Water isotope analyses
Modern water samples from the spring were analyzed for
18
O/
16
O and D/H ratios using a
spectroscopic Liquid-Water Isotope Analyzer (Los Gatos Research, Inc) at the University of
California at Davis or cavity ring-down spectroscopy (Picarro Inc.) at the SIRFER laboratory at
the University of Utah for isotopic determination of δ
18
O and δD. Samples are calibrated against
working standards of known isotopic composition, which allow for normalization to the 2-point
Vienna Standard Mean Ocean Water (VSMOW, 0‰) and Standard Light Antarctic Precipitation
! 12
(SLAP, δ
18
O value −55.5‰, δD value −428‰) isotopic scale. The precision of replicate
injections (1σ) of these standards was better than 0.2‰ and 2‰ for δ
18
O and δD, respectively.
RESULTS
Facies descriptions
We targeted four of the most prominent carbonate exposures along the fluvial path (all
labeled in Fig. 1B and pictured in Fig. 2): (1) perched cascade (34°46!720"N, 120°02!216"W),
(2) fluvial barrage (34°46!643"N, 120°01!219"W), (3) fluvial channel (34°46!040"N,
120°01!281"W), and (4) fluvial cascade (34°46!601"N, 120°01!392"W). The perched deposit is
a succession of carbonates that formed on a steep slope laterally adjacent to the boxed spring but
about 10 m above the present valley floor (Fig. 1B). Carbonates are about 15 m in lateral extent
and up to 2 m thick. The unit terminates with a cascade facies (Fig. 2A) that contains
overhanging carbonate curtains (sensu Pedley, 1990). The carbonate samples investigated in this
study originate from the outermost face of the cascade wall (Fig. 2A).
The barrage carbonates (sensu Pedley, 1990) are the first extensive carbonates (~20 m
lateral extent) that formed on the valley floor about 60 m downslope and about 5-10 m lower in
altitude from the boxed spring. Carbonates crop out and occupy a width of about 15 m along the
valley floor and are less than 0.5 m thick (Fig. 2B). Although most of the carbonate surface is
covered in dense mosses and foliage (Fig. 2B), some laminated structures can be seen in outcrop,
as well as molds of encrusted leaves and plant debris. Samples analyzed from the barrage facies
originate from the top 20 cm of small dams.
The next prominent carbonate facies is a fluvial channel that occurs about 90 meters
down slope of the boxed spring on the north side of the valley (Fig. 2C). The channel is about 2
m wide and about 1 m thick and exhibits successive draping carbonate fabrics. Presently, the
! 13
carbonate surface is covered in thin colonies of dry mosses. Erosive features are common below
the channel where fast flowing water could have scoured and transported carbonate downstream.
Also, successive carbonate drapes could indicate the deposit underwent periods of non-
deposition and possibly erosion. Carbonate samples examined from this deposit originate from
the outermost surface of the terrace face (~20 cm-thick).
The distal end of the fluvial carbonate transect (~about 180 m from the boxed spring)
terminates with a cascade deposit that is about 15 m wide and contains a drop in elevation of
about 10 m from the top to the bottom of the cascade (Fig. 2D). Within the fluvial cascade, at
least two distinct episodes of carbonate growth can be observed on the surface of the deposit: (1)
a basal white to light brown deposit widely distributed on the cascade face and (2) a brown,
highly porous carbonate with a patchy distribution, typically restricted to small (~10 cm wide
and ~5 cm thick) dams that drape the basal white deposit. The total thickness of the cascade
carbonate is about 0.5 m. The carbonate samples examined from this deposit originate from the
basal white to light brown carbonates and were collected near the top of the cascade (upper ~20
cm).
All of the carbonate facies examined contain carbonate pieces that are not attached to the
deposit potentially indicating they were transported from up the valley. In particular, large (m-
scale) talus blocks eroded from the perched cascade and lie unconformably along the valley
channel.
Mesostructure
Representative polished rock samples from the four facies are pictured in Fig. 3.
Mesoscopically, the samples contain bands of variable thicknesses that range from about 1mm to
5 mm. Polished samples from the perched deposit are largely composed of dense light and dark
! 14
bands. In cross section, the barrage deposits, fluvial channel, and fluvial cascade samples all
contain abundant macro- and microphyte molds (Fig. 3C). Additionally, there is considerable
intra- and inter-fabric variability among the different facies, despite all having been collected
from the top ~10–20 cm surfaces of each deposit. While the internal structure and thicknesses of
bands varies across the different deposits, a prominent similarity is the occurrence of
conspicuous light brown bands (~3 mm to 5 mm thick; see arrows in Fig. 3) that form a marked
contrast with the dominant dark brown color of other fabrics. The bands are composed of
isopachous fibrous calcite crystals (Fig. 4). In some cases, the calcite grew radially, likely
around twigs or other debris that were deposited in the spring system. Its fibrous nature is better
observed in a freshly broken piece (Fig. 4A). Light brown and dark, dense bands commonly
alternate with one another and are both composed of calcite (Fig. 4B).
Microstructure observations of light brown and dark brown bands
In contrast to the polished hand samples, the dark, dense bands appear lighter in color in
transmitted light and the light brown fibrous bands appear darker (contrast Fig. 4B with Figs.
4C-E). The dark, dense bands are composed of isopachous bladed calcite crystals with well-
developed pyramidal terminations (Fig. 4D), whereas the light brown bands are composed of
bundles of calcite tubules (Fig. 5E). The tubules bifurcate along their length to form fan-shaped
bundles (Figs. 5A-B). Each fan-shaped tubule bundle is approximately 1.5 mm in length and
averages about 0.5 mm in thickness (see outline of a fan-shaped tubule bundle in Fig. 5A). An
oblique SEM view of the bottom of several tubule bundles displays that each bundle originated
from a common point (Fig. 5C). Figs. 5D-E illustrate the initially hollow nature of the tubules.
The tubules have a consistent diameter in vertical cross section (20 ± 2 µm). The consistent
diameter of the tubules is further confirmed via a horizontal cross-section of a tubule bundle
! 15
(Figs. 5F-G). Under cross-polarized light, the tubule walls in a given bundle exhibit unit
extinction (Fig. 5B) indicating each bundle of tubules is a monocrystal.
Carbonate isotopic analyses
The oxygen and carbon isotopic values of micro-drilled carbonate samples are listed in
Table 1, differentiating their fabric type: light brown bands (LBB) and dark brown bands (DBB).
Samples marked with an asterisk in Table 1 indicate the mean δ
18
O and δ
13
C values of at least
three measurements along the respective sample band. The standard deviation (1σ) along a
single band is less than 0.70‰ and 0.20‰ for δ
13
C and δ
18
O, respectively.
The δ
18
O values of the LBB from all four depositional facies range from –6.78‰ to –7.48‰
with a mean value of –7.12 ± 0.20‰ (1σ, n = 14). The DBB exhibit δ
18
O values that range from
–6.77 to –7.66‰ with a mean value of –7.27 ± 0.28‰ (1σ, n = 13). The mean δ
13
C value for the
LBB is –8.92 ± 0.65‰ (1σ, n = 14) and the DBB exhibit a mean δ
13
C value of –9.17 ± 0.37‰
(1σ, n = 13).
Water isotopic analyses
Water samples from the boxed spring were sampled eight times between 2009 and 2012.
Measured δ
18
O and δD values are listed in Table 2 with their corresponding collection date. The
isotopic composition of modern spring water is relatively constant with a mean δ
18
O value of –
7.58 ± 0.23‰ (1σ, n = 8) and δD value of –46.7± 2.3‰ (1σ, n = 8). Paired δ
18
O and δD values
place the modern spring water as slightly enriched above the global meteoric water line.
DISCUSSION
Oocardium stratum calcite biosignature
The fan-shaped tubular crystal structure composed of 20 µm diameter bifurcating tubules
is strikingly similar to the unique calcite microstructure formed in the presence of the desmid
! 16
microalgae Oocardium stratum (Nägeli, 1849). Oocardium is a colonial desmid whose cell
diameter ranges from 17–20 µm and is known to live in colonies of about 100 cells (Pentecost,
1991). Individual cells release copious amounts of mucilage preferentially on one end of the
cell. Calcite nucleates on and around the mucilage portion and as the cell divides, the calcite
takes the shape of the secreted mucilage tube. Continued cell division results in a monocrystal
fan-shape bundle composed of bifurcating calcite tubules, see Golubić et al. (1993) and Sanders
and Rott (2009) for growth illustrations. Live cells remain above the calcite rim via continuous
secretion of mucilage, which is believed to help elevate the cell (Golubić et al., 1993). Former
calcite colonies of Oocardium calcite (labeled ‘c’ in Fig. 6A) can be identified in thin section
based on the presence of the tubules, their consistent diameter, bifurcation pattern, and unit
extinction pattern of the tubule walls in cross-polarized light highlighting a monocrystal (Fig.
5B). The boundary of individual fan-shaped bundles is well-defined in thin section (Figs. 6B-C).
A horizontally and vertically oriented boundary between two fans displays the empty tubules that
likely once enclosed an Oocardium cell (Fig. 6D).
Environmental conditions associated with Oocardium calcite deposition
Oocardium stratum is known as an unusual desmid for the way it becomes encrusted in
calcite and for being endemic to freshwater carbonate depositing sites (Wallner, 1933; Pentecost,
1991). In sites that contain extensive active Oocardium colonies, a water temperature range of
~9–13 °C is frequently reported (Rott et al., 2009; Sanders and Rott, 2009; Gesierich and Kofler,
2010; Linhart, 2011), suggesting the Zaca carbonate fabrics were sourced from cold water rather
than thermal waters. Oocardium is known to exist in slightly warmer or colder conditions
(Pentecost, 1991; Golubić et al., 1993; Gradzinski, 2010); however its growth was not observed
throughout the year and the possibility of water temperature controlling its development and
! 17
persistence was not directly addressed. Linhart (2011) noted that optimal growth conditions for
Oocardium appeared at a water temperature of 13 °C based on a 17 month study on actively
growing Oocardium colonies that were monitored on a weekly basis.
Another common growth element of Oocardium stratum is its preference for swiftly to
fast flowing water (Mathews et al., 1965; Pentecost, 1991). Although Gradzinski (2010) reports
growth under ‘sluggish flow’, growth in water that accumulates in pool facies has not previously
been observed. This characteristic is consistent with reports of Oocardium typically occurring in
association with cascade/waterfall facies (Pfiester, 1976; Golubić et al., 1993; Pentecost and
Zhang, 2000; Rott et al., 2009), i.e., areas that when wet will always experience flow. The
samples for this study were collected from the cascade facies, further corroborating our
interpretation of the structures. The occurrence of Oocardium in the barrage facies indicate they
could have formed during an episode of very fast flow accumulating of the edges of small dams,
where our samples originate from. Given the intimate relationship between Oocardium and
calcite deposition, it is reasonable that Oocardium colonies would prefer to colonize sites that
degas CO
2
quickly and consequently favor rapid carbonate deposition.
Considering the spring was boxed in 1911, we assess whether carbonate deposition
would occur today if the spring orifice had been left undisturbed, in order to establish whether or
not the carbonates represent conditions wetter than today. Measurements of modern spring flow
rates (out of the ~3 cm diameter pipe attached to the spring box) are on the order of 0.3 L/s after
dry years and 0.7 L/s after wet years (Norris and Norris, 1994). Sanders and Rott (2009) report
values of 1–2 L/s and 5–10 L/s from two different sites in the Austrian Alps where Oocardium
grows today. Comparing these Alpine values to those from Zaca suggests that despite spring
capture, the flow rates emerging from the spring today are insufficiently high to sustain actively
! 18
growing Oocardium colonies. Furthermore, the occurrence of Oocardium at the distal cascade
facies (Fig. 2D) implies that flow rates at the spring orifice would have had to be sufficiently
high to maintain the Alpine Oocardium threshold rates (at least 1 L/s) throughout the extent of
lateral transport (~180 m) when these waters reached the cascade. The presence of the
Oocardium bearing tufa at the distal cascade facies thus suggests a significantly wetter climate
regime during its deposition.
Oxygen Isotope Thermometry
We utilize the Hays and Grossman (1991) oxygen isotope paleotemperature equation:
TºC = 15.7 – 4.36(δ
18
O
calcite
− δ
18
O
water
) + 0.12(δ
18
O
calcite
− δ
18
O
water
)
2
and compare values to
temperature estimates derived using the Kim and O’Neil (1997) equation re-expressed by Leng
and Marshall (2004) as: TºC = 13.8 – 4.58(δ
18
O
calcite
− δ
18
O
water
) + 0.08(δ
18
O
calcite
− δ
18
O
water
)
2
to
obtain a range of depositional temperature estimates (Table 1). These equations were derived for
calcite (see review in Grossman, 2012) and are thus applicable to our study, as calcite was
determined to be the sole mineralogy. Both of these equations are frequently applied to
freshwater carbonates (Garnett et al., 2004; Makhnach et al., 2004; Pentecost, 2005; Andrews,
2006). Differences in temperature estimates between these two equations are empirical, where
temperature estimates utilizing the Hays and Grossman (1991) equation, yield warmer
temperatures than the Kim and O’Neil (1997) equation (+1.8º and +3.3 ºC at 25º and 0º
respectively) (Grossman, 2012).
Key assumptions inherent in these temperature calculations are that (1) the carbonate
precipitated in isotopic equilibrium with the spring water; (2) the water temperature during
deposition as well as (3) the isotopic composition of spring water were not substantially different
from today during carbonate formation. We now explore to what extent these assumptions were
! 19
problematic. For carbonate precipitation that occurs near the spring orifice, assumption (1) is
often incorrect as rapid degassing might cause carbonates to precipitate before the spring water
equilibrates to local atmospheric conditions (Turi, 1986). In this study, carbonates from the
fluvial deposit were collected starting at about sixty meters from the spring box potentially
reducing the effects of disequilibrium precipitation. Assessing assumption (2), modern water
temperature has been measured twice by inserting a thermometer into the pipe downstream from
the box at 14 ºC and 13 ºC, indicating stable spring water temperatures in a shaded riparian
corridor during different seasons. Since temperatures in sub-tropical regions are thought to have
not fluctuated by more than 1−2 ºC over glacial-interglacial timescales (MARGO, 2009), then
this temperature estimate appears to be a reasonable assumption for the past. Assumption (3) is
that the mean δ
18
O of modern spring water approximates that of past spring water (Table 2). If
the carbonates formed from water with a similar δ
18
O to that of today, with carbonate
precipitation at or near isotopic equilibrium, then the paleotemperature estimates (Table 1) are
similar to modern—in this scenario there was minimal change in either temperature or
precipitation isotopes. However, if the carbonates formed during the last glacial it is possible
that the spring water isotopic composition was more depleted than today. As an upper limit on
how much more depleted we take the ~1.5‰ depletion in mean glacial versus Holocene δ
18
O
values of calcite from Moaning Cave in the Sierra Nevada Mountains of California (Oster et al.,
2009), which is likely to see a larger shift because of its more inland and elevated location. If the
δ
18
O of spring water was 1‰ more depleted at the time of formation then this would suggest
temperatures of 9.9 ± 1.0 ºC (using the equation of Hays and Grossman, 1991) and 7.6 ± 1.1 ºC
(using the equation of Kim and O’Neil, 1997), which appear to be too cold to be reasonable
given the limited expected range of sub-tropical, coastal terrestrial temperature changes. Thus
! 20
we conclude that a scenario of minimal change in the δ
18
O values of spring water and
temperature is most likely.
The mean temperature estimate using the Hays and Grossman (1991) equation is 14.1 ±
1.1 ºC and 12.1 ± 1.1 ºC using the Kim and O’Neil (1997) equation (Table 1). Irrespective of
these modest differences between the two equations, modern temperature observations (14 ºC
and 13 ºC) are within the calculated ranges from each equation (Fig. 7). Overall, the calculated
temperatures correspond to spring water temperatures normally associated with non-thermal
waters (Pentecost, 2005; Cepezzuoli et al., 2014), suggesting that the carbonates formed from
ambient temperature. However, it is important to note that deposits sourced from cooled, deeply
cycled (geothermal) waters have been previously described (e.g., Zhang et al., 2012),
highlighting the importance of integrating textural and geochemical data to deposits of unknown
affinity. Here, ambient temperature estimates (Table 1) along with the presence of Oocardium
discredit the possibility of geothermal influence. Furthermore, the similarity of the δ
18
O of the
carbonate compared to modern water δ
18
O (Table 2), strongly suggests the carbonate precipitated
from groundwater recharged from meteoric waters, under temperatures and source water isotopic
compositions relatively similar to modern. The observed variability in δ
18
O of the carbonate
within the deposits is more likely due to changes in water temperature than to source water
isotopic changes given the residence time in the groundwater is expected to even out such
variability.
Temperature estimates for the LBB are lower on average compared to the DBB (Fig. 7),
possibly indicating that water temperature is a controlling factor on the development of
Oocardium. However, other factors like flow rates, nutrients, water turbulence etc., might also
play a significant role on the development of Oocardium, supported by the observation that
! 21
although common, the Oocardium fabric is not the only crystal fabric type (Fig. 3). Several
studies of cold-water carbonates have demonstrated the potential for annual or seasonal banding,
offering the possibility for very high-resolution paleoenvironmental analyses (Matsuoka et al.,
2001; Kano et al., 2004; Andrews and Brasier, 2005). Bands associated with active Oocardium
deposition have also been shown to be seasonal (Sanders and Rott, 2009), where Oocardium
growth rates were greatest in the spring and summer. Presently, >80% of precipitation in Santa
Barbara occurs in the winter (Cayan and Rhoads, 1984) suggesting that deposition may only
have occurred in the winter and that δ
18
O of winter rainfall would be a key parameter controlling
the δ
18
O of calcite.
Carbon Isotopes
Most analyses of ancient spring-associated carbonates utilize stable isotope analyses to
determine the origin of the source waters (Andrews, 2006). Our carbon isotope values are
strongly indicative of soil-derived CO
2
with an overall δ
13
C range of –9.92 to –7.01‰. The wide
range in δ
13
C and a narrower range in δ
18
O is commonly reported in other fluvial carbonate
systems (e.g., Chafetz et al., 1991). Low covariance between δ
13
C and δ
18
O (Fig. 8A) may be
attributed to a low residence time of water in the aquifer resulting in minimal evaporation as is
indicated by only a slight enrichment of δ
18
O modern spring values compared to local
precipitation and relative to the global meteoric water line (Feakins et al., 2014). Such a low
degree of covariance is typical of hydrologically open systems and in this case, the δ
18
O
composition reflects local meteoric inflow (Feakins et al., 2014). Considered together, the δ
13
C
and δ
18
O of the carbonate along with the estimated depositional temperatures shown in Table 1,
present strong support for meteoric recharge of the groundwater aquifer resulting in spring water
that did not undergo significant evaporation.
! 22
Given the intimate association of Oocardium with calcite deposition, we plot δ
13
C values
of light brown bands versus the dark brown bands in order to investigate the potential for
biological photosynthetic fractionation of the resulting δ
13
C (Fig. 8B). Our results reveal a
significant δ
13
C enrichment (t = 2.86, p<0.05) in the Oocardium fabric compared to the dark
brown bands. To date, the degree to which microorganisms can impart a photosynthetic effect
on δ
13
C of calcite via preferential uptake of
12
CO
2
remains to be resolved, as it is usually
complicated by the wide range of factors that may control the δ
13
C value of calcite (Andrews,
2006). Some studies have shown a potential for biological fractionation of
13
C/
12
C relative to
bulk calcite involving cyanobacteria and some algae (Pentecost and Spiro, 1990; Arp et al.,
2001), while others have shown that this discrimination is not always detectable in the calcite
δ
13
C (Shiriashi et al., 2008). Garnett et al., (2004) found that differences in δ
13
C could be related
to seasonality where during wet seasons δ
13
C reflects relatively lighter values derived from
‘light’ soil zone carbon whereas during drier seasons the δ
13
C contains a stronger relatively
‘heavy’ bedrock signature. However, the relatively enriched δ
13
C values we observe correspond
to Oocardium calcite whose depositional conditions are dependent upon high flow/moisture,
thereby instead providing potential support for photosynthetic fractionation rather than
bedrock/soil zone differences. The present study is the first to report calcite C and O stable
isotope values associated with Oocardium deposition, but future work specifically on actively
growing colonies of Oocardium is required to better constrain and assess biotic fractionation
associated Oocardium calcite formation.
To investigate any lateral variability across facies, we plot the carbonate δ
18
O and δ
13
C
against the distance downstream from the modern boxed spring (Fig. 9). Assuming all of the
water supplying carbonate growth once originated from waters similar to the boxed spring site,
! 23
we might expect a slight increase in depositional temperature downstream and/or an overall
enrichment of isotopic values resulting from progressive evaporation and CO
2
degassing (Turi,
1986; Pentecost and Spiro, 1990; Matsuoka et al., 2001). However, our data do not appear to
record a pattern of (1) enrichments of δ
18
O or δ
13
C and thereby (2) enrichments in calculated
temperature with progressive distance from the modern spring orifice (Fig. 9). Thus, it is
possible that there is slight temperature variability each time each facies formed or that certain
facies experienced more/less evaporation. Additionally, variables like changing flow rates and
flow paths during deposition of different facies, changes in relative soil/plant/biologic
contribution, and the potential for additional underground water recharge downstream may cause
the stable isotope values to lack a facies trend with distance from the spring orifice (Arenas-
Abad et al., 2010).
Significance for climate studies in southern California
We have established that inactive spring carbonates near Zaca Lake in Santa Barbara
California were sourced from ambient temperature water with a strong soil zone δ
13
C signal
(Table 1). The presence of Oocardium constrains its associated depositional environment
including spring flow, temperature, and humidity. Based on studies of European tufa, Pentecost
(1995) noted that tufa deposits in general occur in areas characterized by average annual
temperatures ranging between 5–15 °C and annual rainfall exceeding 500 mm, although tufas
from semi-arid environments have been described. Tufa sites with active Oocardium stratum,
however, are more commonly found where rainfall exceeds 1,000 mm per year. Modern tufa
sites reported in Montgomery County Virginia and the Arbuckle Mountains Oklahoma both
contain extensive Oocardium stratum growth and both of these sites experience more than 1,000
mm of rainfall a year. Additionally, a study by Gradzinski (2010) documented the growth rates
! 24
of four active tufa deposits of Poland and Slovakia. Oocardium growth was only observed at the
site that experiences greater than 1,000 mm of rainfall per year, which suggests that perhaps an
ample water supply is also a controlling factor on Oocardium growth.
Zaca Lake currently experiences an average winter temperature of 9 °C and an average
annual precipitation of ~740 mm (PRISM data, 1895-2011; Daly et al., 2008), with high inter-
annual variability and over 80% of the rainfall occurring between October and March. The
presence of this carbonate deposit and the occurrence of Oocardium stratum constitute
compelling evidence for much wetter conditions in this region during carbonate formation. Of
particular importance is the occurrence of the perched deposit (presently ~10 m above the boxed
spring site) indicating a time of deposition for the perched cascade when the water table was
significantly higher, as discussed in detail in the subsequent chapter.
CONCLUSIONS
This study explores the carbonate fabrics and associated stable carbon and oxygen
isotopic composition of currently-inactive spring carbonate deposits near Zaca Lake in Southern
California. Stable isotopic analyses of carbonate δ
18
O are consistent with deposition in ambient
temperature water (~11−16 °C). A mean δ
13
C value of −9.00 ± 0.62‰ (1σ, n=27), is
characteristic of soil derived CO
2
. The discovery of the calcite microstructure of the unusual
desmid Oocardium stratum across all facies surveyed provides complementary support to the
stable isotopic data indicating that this is a tufa (ambient) rather than a travertine (hot spring)
deposit, highlighting the utility of the Oocardium calcite microstructure as an indicator of
depositional water temperature and flow conditions in ancient/inactive deposits. The presence of
cold-water spring carbonates influenced by groundwater input has potential to record past
! 25
pluvials. We conclude that the carbonates grew at some time in the recent past when conditions
were wetter as discussed further in Chapter 3.
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! 32
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! 33
FIGURES
Figure 1. Geologic setting of spring-associated carbonates. (A) Geologic map of the Zaca Lake
catchment. Abbreviations: M = Monterey; Qs = surface Quaternary; L = landslide; Tv = Tertiary
volcanics. Dashed lines indicate fault lines. Contours are at 200 m intervals. (B) Schematic
representation of carbonate deposit (imitating Viles et al., 2007) with the labeled location of the
boxed spring, fluvial deposit, and spring deposit (the image is not drawn to scale; the
approximate distance from the boxed spring to the cascade is ~180 m).
! 34
Figure 2. Facies of the spring carbonate deposits. (A) Cascade facies of the perched deposit
displaying overhanging carbonate curtains. (B) Barrage deposits. (C) Fluvial channel. (D)
Terminal fluvial cascade deposit with people on the outcrop for scale.
! 35
Figure 3. Mesostructure of the spring carbonate fabrics from the four facies in Fig. 2. (A) Banded
carbonate from the outermost cascade wall of the perched deposit. (B) Broken carbonate piece
from the fluvial terrace channel. (C-D) Microphyte-encrusted carbonate and banded carbonate
from the barrage deposits. (E) Banded carbonate from the terminal cascade facies. White arrows
denote the light brown bands.
! 36
Figure 4. Meso and microstructure of the spring carbonate fabric. (A) Freshly broken piece
displaying the texture of the light brown bands (LBB). (B) Polished hand sample denoting the
contrast between the dark brown bands (DBB) and LBB. (C) Photomicrograph of a light and
dark couplet. (D) Well-defined calcite crystal terminations comprising the DBB. (E) Calcite
tubules (~20 µm in diameter) comprising the LBB.
! 37
Figure 5. Microstructure of the light brown bands. (A) Several adjacent calcite tubule bundles.
(B) Adjacent calcite tubule bundles (same as Fig. 3A) under cross-polarized light. (C) Oblique
SEM image of the bottom of several tubule bundles demonstrating that each bundle originates
from a common point. (D) Photomicrograph of calcite tubules. (E) SEM image of the top of
several tubules displaying their hollow nature. (F) Horizontal cross section through a tubule
bundle showing round voids of consistent diameter (~20 µm). (G) SEM image displaying tubule
voids.
! 38
Figure 6. Photomicrographs of the Oocardium stratum calcite microstructure. (A) Three distinct
oblique sections of Oocardium colonies labeled with a “c”. (B) Several adjacent tubule bundles.
(C) The bundles are composed of tubules whose growth terminates at the intersection with other
fans. (D) Horizontally and vertically oriented junction of two tubule bundles, arrows denote sites
that Oocardium cells might have once inhabited.
! 39
Figure 7. Box and whisker plots of calculated temperature estimates comparing the light brown
and dark brown bands, where the whiskers represent the range, the box represents the upper and
lower quartiles and the black horizontal line within the boxes denotes the median.
Figure 8. Carbonate stable carbon and oxygen isotope results. (A) Carbonate δ
13
C and δ
18
O cross
plot. (B) δ
13
C box and whisker plots (see caption of Fig. 7) of light brown bands and dark brown
bands.
! 40
Figure 9. Stable isotope plots versus distance from the boxed spring. (A) δ
18
O versus distance
from the boxed spring. (B) δ
13
C versus distance from the boxed spring. Note: Data from the
perched cascade is not included in these plots as those carbonates formed away from the boxed
spring fluvial path.
! 41
Table 1. Stable isotopic compositions of δ
18
O and δ
13
C (‰ VPDB) for carbonate from their
corresponding facies
*Samples whose values represent the mean δ
18
O, δ
13
C, and temperature of at least three measurements along the
respective sample band. LBB: light brown bands; DBB: dark brown bands
Facies and
Distance
from
boxed
spring (m)
Fabric
type
Sample
δ
13
C
carb
δ
18
O
carb
δ
13
C
carb
1σ (‰)
δ
18
O
carb
1σ (‰)
Calculated
Temp.
(°C)
Hays &
Grossman,
1991
Calculated
Temp.
(°C)
Kim &
O’Neil,
1997
Perched
Cascade
(~10m
above
boxed
spring)
LBB
Zpc-ooc2 -7.01 -6.78 12.2 10.2
Zpc-ooc3 -7.97 -7.41 14.9 13.0
PC2A -8.02 -6.96 12.9 11.0
DBB
Zpc-d1 -8.51 -6.77 12.2 10.2
Zpc-d2 -8.99 -7.23 14.1 12.2
PC2d -9.24 -7.04 13.3 11.4
Fluvial
Barrage
(~60)
LBB
ZPooc13* -9.62 -7.48 0.01 0.13 15.2 13.4
Zb-oc-b -9.21 -7.11 13.6 11.7
Zb-oc-t -8.98 -7.14 13.7 11.8
DBB
ZPd13* -9.92 -7.59 0.13 0.19 15.7 13.9
Zb-d -9.12 -7.48 15.2 13.4
Zpd4 -9.60 -7.06 13.4 11.4
Fluvial
Terrace
Channel
(~90)
LBB
Ztooc12* -8.80 -7.10 0.71 0.20 13.5 11.6
Ztooc45* -8.73 -7.03 0.20 0.08 13.3 11.3
Ztooc6 -9.00 -7.10 13.6 11.6
DBB Ztd12* -9.70 -6.82 0.21 0.01 12.4 10.4
Ztd45* -9.04 -7.33 0.14 0.07 14.6 12.7
Fluvial
Cascade
(~180)
LBB
ZOC2* -9.30 -7.41 0.53 0.03 14.9 13.0
OC -9.24 -6.89 12.7 10.7
FC-01 -8.70 -7.18 13.9 12.0
FC-02 -8.63 -7.16 13.8 11.9
FC03* -8.83 -6.96 0.17 0.06 13.0 11.0
DBB
ZA1 -9.65 -7.35 14.6 12.8
ZM1 -9.16 -7.51 15.3 13.5
ZA2 -9.57 -7.46 15.1 13.3
FC-D1-1 -9.26 -7.66 16.0 14.2
FC-D3* -9.38 -7.30 0.22 0.06 14.4 12.5
Mean Overall -9.01 -7.20 14.1 12.1
1σ 0.62 0.25 1.1 1.1
Mean LBB -8.71 -7.12 13.8 11.8
Mean DBB -9.31 -7.27 14.2 12.2
! 42
Table 2. Stable isotopic compositions of δ
18
O and δD of modern water samples from Zaca Spring
Collection Date Water δ
18
O (‰ VSMOW) δD (‰ VSMOW)
14 Feb 2009 -7.30 -47.0
30 Oct 2009
25 Feb 2010
-7.70 -47.0
-7.70 -46.3
11 Jun 2010 -7.70 -47.0
5 Nov 2010 -7.70 -48.0
5 Apr 2011 -7.80 -48.0
27 Sep 2011 -7.14 -41.4
16 Apr 2012 -7.62 -48.7
Mean -7.58 -46.7
1σ 0.23 2.3
! 43
Chapter 3:
Fluvial tufa evidence of Late Pleistocene wet intervals from
Santa Barbara, California, U.S.A
ABSTRACT
Past pluvials in the western United States provide valuable context for understanding
regional hydroclimate variability. Here we report evidence of conditions substantially wetter
than today from fluvial tufa deposits located near Zaca Lake, Santa Barbara County, California
that have been dated by radiocarbon (
14
C) and Infra-Red Stimulated Luminescence (IRSL). Two
successions of tufa deposition occur within a small catchment that drains Miocene Monterey
Formation bedrock: 1) a fluvial deposit (0−0.5 m thick, 200 m in extent) that formed along a
narrow valley below a modern spring, and 2) a perched deposit about 10 m higher (2 m thick, 15
m in extent). IRSL and radiocarbon dating of the perched carbonates suggests at least two
episodes of carbonate growth: one at 19.4 ± 2.4 (1σ) through 17.8 ± 2.8 (1σ) ka and another at
11.9 ± 1.5 (1σ) ka verified with a charcoal
14
C age of 10.95 ± 0.12 (2σ) cal ka BP. The
relationship between the perched and fluvial spring deposits is inferred to represent a drop in the
water table of more than 10 m associated with a transition from a wet climate in the late glacial
to a dry Holocene today.
The wet period indicated by tufa growth between 19.4 and 17.8 ka is relatively consistent
with other California climate records both north and south of Zaca Lake. However, tufa growth
ca. 12 to 11 ka demonstrates wet conditions occurred as far south as Zaca Lake during the
Younger Dryas event, in contrast to climate records farther south in Lake Elsinore indicating
persistently dry conditions through this interval. A small shift north in the average position of
the winter season storm track could explain wet winters at Zaca while at the same time
! 44
generating dry winters at Lake Elsinore, 275 km southwest of Zaca. If true, these data indicate
that rather small latitudinal shifts in the average winter season storm track can produce large
changes in regional hydroclimate.
INTRODUCTION
With concern over the projected aridification of the southwestern US (Seager et al., 2007,
Williams et al., 2013), there is heightened interest in characterizing past evidence for extended
pluvial periods and droughts in California’s pre-instrumental record to better understand the
large scale controls on climate and how these may change in the future. The western United
States (western US) experienced a wetter climate during the Last Glacial Maximum (18-20 ka cf.
Denton et al., 2010) relative to present as evidenced by palaeolakes Bonneville and Lahontan
(Benson, 1990), as well as expanded palaeolakes Estancia and Mojave (Allen and Anderson,
2000; Anderson et al., 2002; Wells et al., 2003). Additional evidence for wetter conditions
comes from elevated and expanded palaeolake shorelines in Owens Valley (Mensing, 2001),
higher sand contribution to profundal sediments in Lake Elsinore (Kirby et al., 2013), isotopic
evidence from speleothems in New Mexico and Arizona (Asmerom et al., 2010; Wagner et al.,
2010), and evidence for vegetation change based on (1) plant leaf waxes from Lake Elsinore
sediments (Kirby et al., 2013) and (2) pollen from marine sediments in the Santa Barbara Basin
(Heusser and Sirocko, 1997). Detailed analyses of those proxy records with high temporal
resolution and improved dating precision have revealed that temperature and hydroclimate were
spatially and temporally variable across the western US in the late glacial and across the
deglacial (Lyle et al., 2012; Kirby et al., 2013). In particular, pollen records from marine
sediments offshore southern and northern California and Oregon reveal variable timing of
hydroclimatic change along the coast (Lyle et al., 2010; 2012). Adding records that concentrate
! 45
on the spatial and temporal picture are key to resolve the history and causes of hydroclimatic
variability for specific regions across the western US (e.g., coastal versus inland). Presently
there are few terrestrial records from coastal southern California that address hydrological
balance at the LGM and across the last deglacial (e.g., Heusser and Sirocko, 1997; Kirby et al.,
2013).
Tufa evidence for pluvial
Spring-associated carbonates generated from carbonate-rich, ambient temperature
groundwater, referred to hereafter as ‘tufa’ sensu Pedley (1990), serve as potential archives of
source waters and climate of their local region (Andrews, 2006). Notably, the presence of large
tufa accumulations in arid and semi-arid regions is indicative of periods of accelerated
groundwater recharge, as carbonate will only form when there is a net recharge to the
groundwater aquifer (Pedley, 1990). Tufa deposits are therefore robust indicators of past
pluvials (Szabo, 1990; Crombie et al., 1997; Auler and Smart, 2001; Viles et al., 2007), and may
serve as proxy records of hydrological balance to complement other local proxy records of
palaeoclimate (e.g., Garnett et al., 2004; Dominguez-Villar et al., 2007; Cremaschi et al., 2010).
The western US hosts several tufa deposits, including the well-known towers and
pinnacles from lakes in Nevada and California (Scholl, 1960; Newton and Grossman, 1988;
Benson, 1994a; Li et al., 2008). Climate reconstructions from these lacustrine tufa deposits and
related lake sediments have added support to the idea the Last Glacial Maximum (LGM) in the
Great Basin was wetter than today (e.g., Benson, 1978). However, records of fluvial tufa
deposits are less well-known (e.g., Barnes, 1965) as they are perhaps less conspicuous compared
to their lacustrine counterparts in the present semi-arid climate of southern California.
! 46
Here we present compelling terrestrial evidence of persistent wet conditions based on
fossil and recent fluvial tufa deposits from a coastal site near Zaca Lake, in Santa Barbara
County, California (Fig. 1). Radiocarbon and IRSL dating allow us to constrain the age of these
deposits. We combine geomorphic, textural, petrographic, and geochemical observations to
evaluate the nature of the depositional environment during this wet interval. We compare these
new findings to other terrestrial and marine records towards better understanding of the regional
patterns and the magnitude, timing, and causes of pluvial conditions in coastal southern
California.
GEOLOGIC AND ENVIRONMENTAL SETTING
Two successions of spring-associated carbonate deposits have been described from the
Zaca Lake catchment ~3 km east of Zaca Lake in Santa Barbara county, California (Fig. 2; Ibarra
et al., 2014). One succession occurs along a narrow valley (referred to hereafter as ‘fluvial’
carbonates) and extends discontinuously for about 200 m with an overall drop in elevation of
about 40 m. The other succession occurs perched upon the slope of the north ridge (referred to
here as ‘perched’ carbonates) about 10 m above the fluvial carbonates (Fig. 2). The carbonates
formed within a relatively small catchment, and drape over Miocene Monterey Formation
bedrock (Fig. 1A). About 100 years ago, the spring was boxed and piped for human
consumption (Norris and Norris, 1994) which continues to the present day, such that carbonate
deposition downstream is likely not active along the entirety of the fluvial path (Ibarra et al.,
2014). The residence time of water in the catchment is relatively short as fluctuations in the
water table on the order of years to decades have been observed in historical documents (Norris
and Norris, 1994), and they are associated with known fluctuations in recorded precipitation for
the region (SB Public Works, 2013). The sensitivity of the spring to decadal scale climatic
! 47
fluctuations suggest that the carbonate precipitation associated with the spring may record
decadal or longer variations in precipitation.
Santa Barbara County is characterized by a Mediterranean climate (warm, dry summers
and cool, wet winters). The region receives about 700 mm of rain each year, with >80% of the
precipitation delivered during the winter between October and March (Cayan and Roads, 1984).
Moisture is advected by westerly winds from the Pacific Ocean (Fig. 1). Knowledge of inter-
annual precipitation variability is limited by the short instrumental record: rain gauge
measurements in Santa Barbara extend back to the 1860s (SB Public Works, 2013). Proxy
evidence can extend our perspective of hydroclimate in the region and is the only way to capture
evidence for multi-decadal to millennial scale droughts and pluvials in the west (Briggs et al.,
2005; Cook et al., 2004; Mensing et al., 2013). The sediments of Zaca Lake have yielded
records of hydroclimate fluctuations over the last 3,000 years including pluvials lasting decades
to centuries based on leaf wax, pollen, and grain size evidence (Feakins et al., 2014; Dingemans
et al., 2014, Kirby et al, 2014). Tufa deposits within the catchment provide evidence for pluvial
conditions back to ca. 20 kyrs as outlined below.
SPRING CARBONATE FACIES
Fluvial carbonates
Carbonates from the fluvial deposits extend for about 200 m from the location of the
spring box to the bottom of the fluvial cascade. However, their distribution is patchy and they
exhibit a maximum thickness of about 0.5 to 1 m. The change in elevation from the spring box
to the top of the fluvial cascade is about 40 m. The most prominent facies are barrage, narrow
channels, and a terminal fluvial cascade unit (Fig. 2). The terminal cascade facies is about 0.5 m
thick and exhibits a drop in elevation from the top to the bottom of the cascade of about 10 m
! 48
(Fig. 3A). Two distinct carbonate textures drape the surface of the cascade unit (Fig. 3B): (1) a
basal white unit heavily encrusted in detrital molds and organic plant debris (Figs. 3B-D) and (2)
a darker, brown surficial carbonate fabric that drapes the underlying white layer and forms small
(~20 cm) dams across the cascade deposit (Fig. 3B). Carbonates along the fluvial unit are
distinctly banded (Fig. 3E). The bands are microscopically composed of alternating isopachous
sparry and micritic laminae (Fig. 3F). Some bands contain ubiquitous calcite microphyte molds,
including the calcite microstructure of the desmid microalgae Oocardium stratum consistent with
depositional water temperatures of ~11 to 16 °C and corroborated by δ
18
O palaeothermometry
(Ibarra et al., 2014).
Perched carbonates
About ten meters above the present-day spring orifice, perched carbonate cascade
deposits occur at the break of the valley-side slope on the north ridge (Fig. 2), representing a
perched spring line tufa facies (Pedley, 1990). Perched spring line tufa is largely controlled by
the slope and rate of water flow over the deposit. Perched springs give rise to prograding
carbonate cascade facies, most of which built outwards resulting in vertical and lateral growth to
form a prominent apron (see Pedley, 1990). Carbonate will prograde in the direction of water
flow and continue to accumulate until the piece fractures (e.g., Pentecost and Zhang, 2008).
Similarly, at our site, a large ~1 m-thick and ~4 m long detached block is oriented so that the
cascading face rests against the wall and the opposite end plunges towards the valley channel
forming a blind cave with the cascade wall (Fig. 4A). Adjacent to the large dipping piece, the
cascade wall exhibits distinct carbonate curtains (sensu Pedley 1990; Figs. 4B-C). Downslope
from the perched cascade, large talus blocks (up to ~2 m in diameter) lie unconformably along
the side slope and the bottom of the valley channel (Fig. 4D). Carbonate samples from the
! 49
perched cascade are highly indurated. The carbonate texture is typified by mesoscopic (~0.5 to 1
cm diameter) vuggy pore space (Fig. 4F). Microscopically, the fabric is dominantly composed
of microspar, micrite, and dog-toothed spar with an irregular, heterogeneous distribution (Fig.
4G).
METHODS
Sample collection
Samples from the fluvial carbonates were collected from the top ~20 cm of areas along
the flow path that contain prominent carbonate accumulation (areas labeled in Fig. 2). Carbonate
from the perched cascade deposit was not easily removed due to the massive and indurated
nature of the deposit. We utilized a handheld drill with a 3 cm diameter drill bit to collect two
cores. One core (70 cm long) was collected from the perched cascade face drilled horizontally
into the cascade wall. A second ‘dark’ core (63 cm long) was collected behind the large
detached carbonate block and immediately transferred into light-proof bags for subsequent IRSL
(Fig. 5A).
Age control
Radiocarbon
Ten Accelerator Mass Spectrometry (AMS)
14
C dates were obtained from carbonate and
organic fragments collected from both of the carbonate deposits. Three pieces of plant debris
were hand picked from within carbonate pieces from the fluvial cascade (~20 cm depth) for
subsequent Δ
14
C analyses. We also dated two organic pieces (detrital twig and root) collected
near the barrage facies (Fig. 2) that each contained concentrically encrusted carbonate (see Fig.
5B); the associated encrusted carbonate was also analyzed for Δ
14
C. Carbonate pieces of distinct
growth phases from the fluvial cascade were also collected for Δ
14
C analyses (see numbered
! 50
labels in Fig. 3B). We also collected carbonate samples from a large block adjacent to the
detached perched cascade. One of these samples was composed of carbonate-cemented clasts of
shale from the Monterey Formation and included a charcoal clast suitable for Δ
14
C analysis.
Samples were sent to the UC Irvine Keck Laboratory for radiocarbon analysis. Δ
14
C values were
converted to calendar years before present using the CALIB 7.0.1 Program (Stuiver et al., 1998),
and the CALIBomb Program (Reimer et al., 2013).
Infra-Red Stimulated Luminescence (IRSL)
A ~63 cm long, 4 cm diameter ‘dark’ core was drilled from a large (~1.5 m thick, ~3 m
long) detached piece of the perched cascade unit (Fig. 5A) for luminescence dating. The core
was extracted and immediately transferred into light-proof black bags for transport to the
laboratory.
The carbonate core was split into five sections under controlled amber and red laboratory
lighting. Two 10 cm core lengths were cut at each end of the core for dose rate estimation, and
the remaining material cut into three sections of approximately 14 cm in length. The three core
sections for dating were each subsequently placed in 3% HCl to dissolve carbonate, ventilated,
but shielded from all light. Acid was replaced until each section of core had dissolved (up to two
weeks), and no further reaction occurred when fresh acid was added. The residual material was
then treated as a standard sediment sample, incorporating the following steps. The samples were
first wet sieved to isolate the 175 – 200 µm fraction. Potassium feldspars were floated from these
fractions using a lithium metatungstate solution of density 2.58 g cm
-3
. After rinsing and drying,
these samples provided just a few hundred grains.
Luminescence dating was performed using a single grain post-IR IRSL approach, based
on the single aliquot regenerative-dose (SAR) method of Baylaert et al. (2009). Measurements
! 51
were performed using a Risø automated TL-DA-20D reader fitted with a dual laser single grain
attachment. Stimulation was provided by a 150 mW 830 nm IR laser at 90%, and luminescence
signals were detected using an EMI 9235QB photomultiplier tube, fitted with a BG3 and BG39
filter combination, allowing transmission in the blue (340 - 470 nm). To reduce thermal transfer
and contributions from slowly bleaching signals, samples were bleached at raised temperature
using Vishay TSFF 5210 870 nm IR diodes.
The SAR protocol used involves repeated cycles with the following steps: 1)
Regenerative beta dose (0 for the natural cycle), 2) Preheat, 60 s at 250 ºC, 3) IRSL
50
, 2.5 s per
grain at 50 ºC, 4) IRSL
225
, 2.5 s per grain at 225 ºC, 5) Test dose, 9.5 Gy, 6) Preheat, 60 s at 250
ºC, 7) IRSL
50
, 2.5 s per grain at 50 ºC, 8) IRSL
225
, 2.5 s per grain at 225 ºC, 9) Hot bleach, 40 s
IRSL using diodes at 290 ºC. The measurement sequence included SAR cycles comprising the
natural measurement, four regenerative-dose points, a zero dose point to assess thermal transfer,
and a repeat of the first dose point to assess recycling. This approach has provided a number of
age estimates spanning 400−80,000 years consistent with independent age control provided by
14
C and
10
Be (Rhodes, submitted).
Most grains measured provided no significant IRSL signal, but a proportion of grains
displayed strong IRSL decays, with linear or saturating growth with increasing regenerative
dose. A degree of variation between the equivalent dose estimates of different grains in each
sample was observed, interpreted as incomplete zeroing of some grains incorporated as the tufa
was building. For each of the three samples, the minimum group of grains, defined as those
consistent with an overdispersion of 15%, was selected; grains with higher dose values were
rejected from the analysis.
! 52
Dose rate estimation was conducted using ICP-OES for K content, and ICP-MS for U
and Th, using the conversion factors of Adamiec and Aitken (1998). The resulting age estimates
are provided in Table 2. We expect that most of the dose rate contributions are from sediment
grains contained within the tufa, but we recognize that tufa can absorb U from water, and the
possibility exists of U disequilibrium. In this case, the ICP-MS estimate of
238
U will
overestimate the dose rate; the U lacks daughter isotopes lower in the decay series, and this
effect can roughly half the beta dose rate from U and have an even more dramatic effect on the U
gamma dose rate (as most gamma energy is provided by isotopes at the end of the U decay
series). To assess the potential magnitude of this effect, we have also calculated the age estimates
with 50% of the U beta dose rate, and no U gamma contribution. We note that this represents an
extreme condition – in practice some or most of the U dose rate contributions probably come
from sediment grains, and for these contributions, we expect secular equilibrium to exist. The
age estimates assuming extreme U disequilibrium (as described above) range from 1 to 2 ka
older than those presented in Table 2. We consider, therefore, that this effect is likely not
disrupting these age estimates significantly. We note the relatively large measurement
uncertainties, caused by having relatively few sensitive grains contributing, and present the age
estimates without allowing for potential disequilibrium effects.
Carbonate isotopic analyses
Isotopic analyses of carbonate oxygen and carbon were conducted on an Elementar
Americas Inc. (Micromass Ltd) Isoprime stable isotope ratio mass spectrometer (IRMS) with a
multi-prep/carbonate device and dual inlet in the Stott lab at the University of Southern
California. Samples were measured relative to CO
2
reference gas calibrated against the NBS-19
(δ
18
O value +2.20‰, δ
13
C value +1.95‰) carbonate standard, which allows for normalization to
! 53
the 2-point VPDB-LVSEC isotopic scale. The precision of this determination is better than
0.06‰ and 0.04‰ (1σ, n = 20) for δ
18
O and δ
13
C, respectively. A working standard (carbonate,
δ
18
O −1.88‰, δ
13
C value of +2.07‰) monitors precision during the course of the run to 0.07‰
and 0.04‰ (1σ, n = 14) for δ
18
O and δ
13
C, respectively.
RESULTS
Age Control
Radiocarbon
Radiocarbon results are listed in Table 2. The piece of charcoal extracted from the
perched cascade has a calibrated age range of 10,830−11,070 cy BP (2σ). All of the organic
fragments from the fluvial deposit contain excess
14
C indicating a significant presence of
radiocarbon from nuclear weapons testing during the 1960s. Samples with excess
14
C have
estimated age ranges from about 1987 to 2007 (Table 2). The carbonate that encrusted the
organic fragments from the fluvial deposits exhibit ages of 10,272 and 18,478 cy BP (Table 2).
Given that the encrusted organic matter recorded modern ages
14
C dating of the encrusting
carbonate, which revealed much older dates, does not accurately reflect the timing of carbonate
deposition, and is not used here for temporal reconstruction. The
14
C ages on carbonate likely
reflect mixing of soil derived CO
2
with old bedrock carbon and are therefore not viewed as
robust to the uncertainties of fraction of ancient carbon so are not considered meaningful
although we do note they are entirely consistent with values obtained on dating the charcoal and
IRSL (Table 1 and Table 2).
IRSL
Three IRSL measurements were obtained for the extracted core (Table 1). The ‘inner’
piece yields a date of 19.4 ± 2.4 (1σ) ka, the ‘middle’ piece an age of 17.8 ± 2.8 (1σ) ka, and the
! 54
‘outer’ piece an age of 11.9 ± 1.5 (1σ) ka. These measurements of IRSL-based age corroborate
the stratigraphic order we would expect given the direction of flow and growth pattern of
perched tufa (Fig. 5A) and are entirely consistent with the charcoal
14
C age from the same
deposit.
Carbonate stable isotopes
Stable carbon and oxygen isotope values from the fluvial and perched cascade are listed
in Table 3 and plotted in Fig. 6. δ
13
C values from the fluvial deposit range from −9.92‰ to -
8.63‰ and exhibit a mean value of −9.21 ± 0.37‰ (1σ, n = 21). Oxygen isotope values from the
fluvial deposits range from −7.66‰ to −6.82‰ and exhibit a mean value of −7.24 ± 0.24‰ (1σ,
n = 21). The δ
13
C of the perched deposits range from −9.24‰ to −6.61‰ with a mean value of
−7.92 ± 0.74‰ (1σ, n = 19) and the corresponding δ
18
O values range from −7.85‰ to −6.77‰
with a mean of −7.21 ± 0.35‰ (1σ, n = 19).
DISCUSSION
Comparison of perched and fluvial deposits
On the basis of field observations and geomorphology, the (1) substantially thicker, (2)
highly eroded, and (3) elevated nature of the perched deposits suggests they formed during an
earlier depositional regime, when the water table was markedly higher (~10 m higher) than it is
today. Facies contrasts between modern and ancient deposits have been reported in the literature
from other arid and semi-arid environments. In these cases, although spring flow is often active
under modern conditions, locations around the spring vent contain carbonate remnants situated at
elevated positions above active springs indicating a drop in the local water table (e.g., Martin-
Algarra et al., 2003; Crombie et al., 1997; Dominguez-Villar et al., 2011; Filho et al., 2012). The
similarity in morphology of our system compared with those reported elsewhere, suggests a
! 55
significant hydroclimate change from higher rainfall (perched deposits) to the present semi-arid
conditions (fluvial deposits).
In addition to geomorphic differences, the textures of the fluvial and perched deposits are
distinct at the meso- and micro-scale (compare Figs. 3E-F with Figs. 4E-F). The banded
morphology of the fluvial deposits is a primary depositional feature typical of freshwater
carbonates (e.g., Kano et al., 2004; Andrews and Brasier, 2005; Golubić et al., 2008). The bands
reflect calcite deposition associated with microalgae that alternates with pyramidal sparidic
bands (Ibarra et al., 2014). In some cases sub-mm diameter pores reflect molds of decayed
organic material that represent the former presence of the microalgae Oocardium stratum (Ibarra
et al., 2014). The continuous nature of the micritic and sparry bands along with well-preserved
microalgal calcite molds strongly suggests the pore space between the bands is largely primary
(Fig. 3F).
In contrast to the fluvial deposit, the highly indurated texture and vuggy porosity of the
perched carbonates suggests they have experienced significant meteoric dissolution and
cementation. Vuggy porosity (Fig. 4E) is considered secondary porosity usually resulting from
the dissolution of calcareous cements (Tucker and Wright, 1990). The patchy distribution and
dominantly micritic nature of the microfabric are typical features of cements that form in the
vadose zone, the zone of undersaturation above the water table (Tucker and Wright, 1990). The
lack of clear stratigraphic structure at the mesoscale (Fig. 4E) together with microscopic textures
that vary at the micrometer scale (Fig. 4F) indicate the perched deposits have undergone several
episodes of dissolution and precipitation.
Textural comparisons between the perched and fluvial deposits reveal different diagenetic
histories despite proximity. Although diagenesis in tufa deposits is not necessarily only
! 56
dependent on the age of the deposits (Pentecost, 1981), the striking contrast between the samples
that originate from the perched deposit compared to carbonate samples collected along the
fluvial channel together with geomorphic differences strongly suggests the perched carbonates
are older. Furthermore, differences in carbonate δ
13
C between the perched and fluvial deposits
may be diagenetic where relatively higher δ
13
C values of the perched carbonates compared to the
fluvial carbonates may have resulted from progressive dissolution and precipitation caused by
percolating rainwater (Janssen et al., 1999). The lack of a clear difference in δ
18
O between the
two deposits (Fig. 6) may be due to diagenetic alteration caused by spring water values with a
similar δ
18
O value to the water that originally deposited the carbonate (e.g., Andrews and
Brasier, 2005).
It is important to note that despite petrographic evidence of diagenesis in the perched
deposit, we do not expect diagenetic alteration to affect the IRSL analyses. IRSL dating does not
focus on the carbonate itself but rather the grains that are trapped while the carbonate forms.
Therefore, although dissolution and precipitation of cements may remobilize trapped detrital
grains, we expect the mobilization to be minimal and not significantly affect the dating
procedure. The correct expected stratigraphic order of the ages obtained (Table 1) further
supports a minimal effect of diagenesis on the IRSL dating result.
Considered together, the (1) geomorphic, (2) textural, and (3) geochemical observations
described above together with
14
C and IRSL ages constitute compelling evidence that the fluvial
and perched deposits reflect distinct depositional periods. The perched deposits formed when the
water table was substantially higher (by at least 10 m), producing thick cascade deposits that
have since undergone significant erosion (Fig. 4D). In our palaeoenvironmental interpretation
below, we focus on the ages obtained for the perched cascade, as carbonate formation about 10
! 57
m above modern spring outflow is directly indicative of past pluvials. Our dating of the perched
cascade suggests at least two episodes of carbonate growth, one ranging from about 19.4 ± 2.4 to
17.8 ± 2.8 ka and the other at 11.9 ± 1.5 ka (Fig. 7A; Table 1). These pluvials were much wetter
than anything in recent history having formed when the water table was ~10 m higher and likely
persisting over thousands of years.
Comparison to regional evidence for wet conditions ca. 19 ka
Although there is extensive evidence for substantially wetter conditions across the
western US, few of the records available represent coastal settings close to the modern
metropolitan centers. Our dated tufa deposits provide ages of carbonate growth representing
substantially wetter conditions than present at this coastal site. Based on stratigraphic
relationships at the outcrop scale, it is likely that earlier pluvials also existed although the timing
of earlier tufa deposition remains to be determined. We focus on comparisons from relatively
proximal coastal sites given recent investigations that highlight key spatiotemporal differences in
hydroclimate between inland and coastal sites (Lyle et al., 2012; Kirby et al., 2013). Pinus
pollen in the marine sediments of Ocean Drilling Program (ODP) Site 893 in the Santa Barbara
Basin (SB Public Works, 2013) about 50 km south of our study site record several episodes of
Pinus expansion (Fig. 7B) interpreted as wet and/or cold conditions (Heusser and Sirocko, 1997).
Pinus pollen is well transported by wind and likely sourced from trees in a near coast region
(Heusser, 2000), including from the Pinus vegetation in the Zaca Lake catchment, which is
native although today enhanced by Forest Service planting and fire suppression (Norris and
Norris, 1994). The onset of a long-lasting wet event at around 21 ka (Lyle et al., 2012) from the
Pinus record correlates well with our age estimates of carbonate deposition for the two inner
parts of the perched core (Fig. 7A). Additionally, a record of grain size (Kirby et al., 2013; Fig.
! 58
7C) and leaf wax δD from Lake Elsinore located about 275 km southeast of our site provide
corroborating evidence for a wet late glacial overlapping with the timing of tufa deposition at our
site. Tufa deposition strengthens the evidence for pulses of wet conditions during the late glacial
in coastal southern California with at least two intervals (ca. 19 and 12 ka) when sustained wet
conditions supported a water table approximately 10 m higher than modern and substantial
carbonate precipitation.
Comparison to regional evidence for wet conditions ca. 12 ka
The Younger Dryas (YD) is a well-known cold period that interrupted the warming out of
the last glacial in the North Atlantic region (Berger, 1990). Although the cooling effects of the
Younger Dryas may not have been global, glaciers were reported to have advanced in the Sierra
Nevada Mountains approximately coinciding with the YD whether due to cooling or wetter
conditions or both (Phillips et al., 2009), and SSTs in the SBB cooled as the California Current
strengthened (Fig. 7D; Pak et al., 2012). In the interior western US, the YD coincided with a wet
period with abrupt onset and termination (Wagner et al., 2010; Asmerom et al., 2012). In coastal
western US the changes in precipitation during the YD (Fig. 7E) are less clear as a result of
dating uncertainties in a variety of records as reviewed by Kirby et al. (2013). Grain size and
δD
(wax)
data from well-dated, high-resolution records from Lake Elsinore suggest no substantial
increase in precipitation during the YD (Fig. 7C). Rather, the onset of the YD at Lake Elsinore
is characterized by an abrupt onset towards less run-off and no obvious change in storm moisture
source (Kirby et al., 2013). Furthermore, there is no apparent termination to the YD; rather,
conditions continue to remain dry into the Holocene (Kirby et al., 2013). In contrast, the ages of
the tufa deposits in the Zaca Lake catchment from our study indicate wet conditions during the
YD. Two age constraints –based on independent techniques– date inclusions within the perched
! 59
tufa and indicate tufa growth: a radiocarbon date on charcoal (10.95 ± 0.12 cal ka BP) and an
IRSL age estimate (11.9 ± 1.5 ka) on silicate grains (Fig. 7A). Together, these ages provide
compelling evidence for carbonate growth at ca. 12 ka. Carbonate growth of the perched tufa, 10
m higher than the present water table, is necessarily associated with conditions that are wetter
than present.
Local cooling is possible because of a decrease in SSTs in the SBB (Fig., 7D; Pak et al.,
2012). The Pinus pollen record from the SBB also yields evidence for wetter terrestrial
conditions (Fig. 7B; Heusser and Sirocko, 1997; Lyle et al., 2012). Our dated tufa deposit
matches the last extreme wet event in the Pinus record (Fig. 7B). Although at much lower
temporal resolution, our record corroborates the SBB record of terrestrial environments,
demonstrating the presence of a wet interval that supported carbonate precipitation at the Zaca
tufa site and Pinus expansion across the area of Santa Barbara County that supplies pollen to the
SBB. Another explanation for wetter winters during the YD at Zaca Lake may be related to a
shift in average winter season storm tracks. A small shift north in the average position of the
winter season storm track could explain wet winters at Zaca as recorded by the tufa growth and
mesic pollen in the SBB while at the same time generating dry winters with no apparent change
in moisture source at Lake Elsinore, 275 km southeast of Zaca. If true, these data indicate that
rather small latitudinal shifts in the average winter season storm track can produce large changes
in regional hydroclimates.
Two other peak wet events are recorded from pollen records in marine sites offshore
Santa Cruz, CA (ODP site 1018) and at the California-Oregon border (ODP site 1019; Lyle et
al., 2012). At ODP site 1018 a significant decline in the prevalence of herbs and shrubs is
interpreted as the initiation of a prominent wet event in the area at ca. 11 ka (Lyle et al., 2012).
! 60
This data matches well with an extended wet interval observed from Moaning Cave on the
western foothills of the Sierra Nevada that lasted from about 12.4 ka to about 9.6 ka (Oster et al.,
2009). Further, the occurrence of a peak wet event off the coast near the California-Oregon
border at ODP site 1019 based on Alnus (alder) pollen also coincides with our tufa record (Lyle
et al., 2010; 2012). Thus far, comparisons between pollen records in northern and southern
California suggest southern California experienced a peak wet event about 6 ka earlier than
central and northern coastal California (Lyle et al., 2010; Lyle et al., 2012). Our new tufa record,
however, suggests coastal southern California remained substantially wet at ca. 12 ka thus
potentially coinciding with wet events north of our site, and challenging the time-transient wet
event proposed.
Missing wet events not captured in the tufa record
Although tufa growth at the perched cascade coincides with the last wet event recorded in
the SBB Pinus record, there are numerous earlier Pinus events that we have not resolved in the
Zaca tufa record to date. Many of these Pinus events are captured in the Elsinore sand record as
intervals of enhanced run-off. Pinus reached a maximum between 14−17 ka, with additional
brief wet events at about 14.1 and 13.3 ka. In particular, the peak in Pinus event at ca. 16 ka is
not particularly well-represented in the tufa, though we note it is within the age uncertainty. A
possible reason for the discrepancy between Pinus events and tufa growth is that compared to
Pinus, tufa deposition reflects substantially wetter conditions, regardless of changes in
temperature, whereas Pinus events—although reflecting wetter and cooler conditions that
present—may not be as responsive to true wet events as those required for perched tufa
deposition. Additionally, the surficial nature of tufa deposition inherently introduces potential
for episodes of erosion and/or non-deposition (Andrews and Brasier, 2005) possibly resulting
! 61
from spring water flow diversion to other channels around the vent. At the moment however,
with the current age uncertainty and dating resolution available on the tufa, it is difficult to fully
resolve the timing of tufa growth during the wettest interval recorded by offshore marine
sediments (Lyle et al., 2010). It is encouraging, however, that both of our main wet events as
recorded by the tufa (ca. 19 ka and ca. 12 ka) entirely span the observed long-lived wet event
derived from marine records 50 km away (Lyle et al., 2012). Our record thus provides an
important complementary addition to evidence for past pluvials in coastal southern California.
CONCLUSIONS
We present
14
C and IRSL dates of perched and fluvial tufa deposits from a coastal site in
Santa Barbara county, California revealing at least two late glacial pluvials. Geomorphic,
textural, petrographic, and geochemical distinctions between perched and fluvial deposits reflect
a transition from a wetter climate regime that resulted in the formation of the perched deposits to
the present actively-accreting fluvial carbonates (located about 10 m below the perched deposits)
along the modern spring flow path. Luminescence dates indicate the perched tufa cascade was
active from about 19.4 ± 2.4 to about 17.8 ± 2.8 ka. These ages agree with independent records
that also suggest wetter coastal conditions at these times. A second wet event dated to 11.9 ± 1.5
ka based on luminescence and verified with a dated charcoal inclusion to 10.95 ± 0.12 ka BP
indicate the Late Pleistocene early Holocene transition was also wetter than present. Ages for
the perched deposits provide a new record of past pluvial conditions for the region, highlighting
the utility of fluvial tufa as a terrestrial record of water balance.
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Allen, B.D., and Anderson, R.Y., 2000, A continuous, high-resolution record of late Pleistocene
climate variability from the Estancia basin, New Mexico: Geological Society of America
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climate change in Southwestern North America at the glacial termination: Quaternary
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FIGURES
Figure 1. Study site. (A) Geologic map of the study area. Abbreviations: M = Monterey; Qs =
surface Quaternary; L = landslide; Tv = Tertiary volcanics. Dashed lines denote fault lines.
Contours are at 200 m intervals. (B) Regional map denoting locations mentioned in the text.
! 71
Figure 2. Schematic of carbonate facies modified from Viles et al., 2007 (figure not drawn to
scale; the approximate distance from the boxed spring to the fluvial cascade is 180 m).
! 72
Figure 3. Multiscale facies of the fluvial carbonates. (A) Terminal fluvial cascade with people
on the outcrop for scale. (B) Distinct carbonate growth episodes along the fluvial cascade with a
basal white carbonate layer (1) and draping brown carbonate layer (2). (C) Carbonate-encrusted
organic debris and plant molds along the fluvial terminal cascade. (D) Carbonate-encrusted twig
located near the fluvial channel. (E) Banded carbonate from the fluvial cascade. (F)
Photomicrograph of (E), blue areas denote pore space.
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Figure 4. Multiscale facies of the perched carbonates. (A) Dipping carbonate block showing
location of the ‘dark core’. (B) Carbonate curtains of the cascade face, bottom. (C) Carbonate
curtains of the cascade face, top. (D) Carbonate talus blocks along the valley slope, looking
northeast. (E) Mesostructure of the perched carbonates with distinct vuggy pores. (F)
Photomicrograph of (E), blue areas denote pore space.
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Figure 5. Dated samples. (A) Location of ‘dark core’ collected for IRSL. (B) Carbonate-
encrusted root for which carbonate (sample ZC-F-root) and organic carbon (sample ZC-F-carb)
14
C was measured.
! 75
Figure 6. Cross plot of carbonate carbon and oxygen stable isotope values of the fluvial and
perched deposits.
! 76
Figure 7. Regional comparisons. (A) IRSL measurements from this study and
14
C of charcoal.
(B) Pinus pollen record from ODP site 893 in the Santa Barbara Basin (Heusser and Sirocko,
1997). (C) Record of percent sand from Lake Elsinore (Kirby et al., 2013). (D) Ca/Mg record
from ODP site 1017E near the Santa Barbara Basin (Pak et al., 2012). (E) NGRIP δ
18
O record
from Greenland highlighting the Last Glacial Maximum (LGM), Older Dryas (OD), Bolling-
Allerod (BA), and Younger Dryas (YD) intervals (North Greenland Ice Core Project members,
2004).
! 77
Table 1. IRSL measurement values of carbonate samples from the perched cascade core.
Table 2.
14
C age data for carbonate and organics extracted from the fluvial and perched deposits
a
Calibrated ages on CALIB 7.0.1 Program (Stuvier et al., 1998)
b
CaliBomb Program (Reimer et al., 2013)
Lab Code D
e
(Gy) ± 1 s Dose rate (mGy/a) ± 1σ Age
(years before AD 2014)
J0311 11.8 ± 1.2 0.99 ± 0.07 11,900 ± 1,500
J0312 18.3 ± 2.5 1.02 ± 0.07 18,000 ± 2,800
J0313 20.3 ± 2.1 1.05 ± 0.07 19,400 ± 2,400
UCIAMS Sample ID
14
C age
(BP)
± 2σ Range
a
Age (cy BP)
a
114961 ZC-carb 17,910 50 21,500−21,899 21,700
124439 ZC-C-carb 15,210 40 18,343−18,612 18,478
124440 ZC-F-carb 9,155 25 10,239−10,304 10,272
124441 ZC-FC-carb 9,960 25 11,263−11,410 11,337
124430 ZC-PC-char 9,600 25 10,830−11,070 10,953
Samples with modern
14
C F (
14
C) ± Year AD
b
115269 ZC-twig 1.1703 0.0023 1987.9
115270 ZC-OM1 1.1919 0.0026 1988.6
115271 ZC-OM2 1.0909 0.0021 1999.9
124431 ZC-C-twig 1.1785 0.0023 1988
124432 ZC-F-root 1.0698 0.0021 2007
! 78
Table 3. Stable isotopic compositions of δ
18
O and δ
13
C (‰ VPDB) for carbonates from their
corresponding deposit.
*Samples whose values represent the mean δ
18
O and δ
13
C of at least three measurements along
the same band.
Carbonate
deposit
Sample ID δ
13
C δ
18
O
Fluvial ZPooc13* -9.62 -7.48
Zb-oc-b -9.21 -7.11
Zb-oc-t -8.98 -7.14
ZPd13* -9.92 -7.59
Zb-d -9.12 -7.48
Zpd4 -9.60 -7.06
Ztooc12* -8.80 -7.10
Ztooc45* -8.73 -7.03
Ztooc6 -9.00 -7.10
Ztd12* -9.70 -6.82
Ztd45* -9.04 -7.33
ZOC2* -9.30 -7.41
OC -9.24 -6.89
FC-01 -8.70 -7.18
FC-02 -8.63 -7.16
FC03* -8.83 -6.96
ZA1 -9.65 -7.35
ZM1 -9.16 -7.51
ZA2 -9.57 -7.46
FC-D1-1 -9.26 -7.66
FC-D3* -9.38 -7.30
mean -9.21 -7.24
1σ 0.37 0.24
Perched
Zpc-ooc2 -7.01 -6.78
Zpc-ooc3 -7.97 -7.41
PC2A -8.02 -6.96
Zpc-d1 -8.51 -6.77
Zpc-d2 -8.99 -7.23
PC2d -9.24 -7.04
PC2A -8.02 -6.96
PC2F -7.98 -7.04
PC2G -8.01 -6.81
PC2D -9.24 -7.04
PC2E -8.16 -6.82
PC5A -7.25 -7.15
PC5B -8.03 -7.85
PC5C -7.91 -7.55
PC5D -8.09 -7.51
PC9A -7.27 -7.63
PC9B -7.19 -7.71
PC9C -6.61 -7.56
PC9D -6.97 -7.43
mean -7.92 -7.21
1σ 0.74 0.35
! 79
Chapter 4:
Microfacies of the Cotham Marble: A tubestone carbonate
microbialite from the Upper Triassic, Southwestern United
Kingdom
ABSTRACT
A remarkably aerially extensive (~2,000 km
2
) unit of carbonate microbialites occurs in
many Triassic-Jurassic boundary interval outcrops of the southwestern United Kingdom and
captures petrographic evidence that could link them to the End-Triassic extinction event. The
bioherms—known regionally as the Cotham Marble—occur as discrete ~20 cm thick, decimeter
to meter-scale mounds, and display at least five growth phases that alternate between laminated
and dendritic mesofabrics. Cross-sections parallel to bedding through the dendritic phases
expose a reticulate dendritic framework separated by polygonal spaces (~1–3 cm diameter),
characteristic of “tubestone” microbialites. Microscopically, the dendrolites contain evenly
distributed rod to filamentous putative microfossils (~2 µm diameter and ~10 µm in length) in a
matrix of micrite, and contain higher TOC than the surrounding matrix. Round to ellipsoidal
spar-filled regions (~200 µm in diameter) within the dendrolites (previously interpreted as
serpulid worm tubes) likely resulted from the production of gas bubbles within rapidly lithifying
mats or are a two-dimensional artifact of evenly spaced three-dimensional branching within the
mats. The fill between the dendrolites of the first layer contains abundant phycoma clusters of
the green algal prasinophyte Tasmanites, commonly considered a “disaster taxon”. The cyclic
phases represent episodic and laterally extensive environmental change within shallow water
coastal environments during a marine transgression. Collectively, the presence of microbial
micrite in a shallow marine setting, the marked lateral extent of the bioherms, and the abundance
! 80
of Tasmanites suggest the Cotham Marble microbialites formed during the high pCO
2
and
relatively warmer conditions associated with the events of the End-Triassic mass extinction.
INTRODUCTION
The Upper Triassic beds of the southwestern United Kingdom (SW UK) contain a
laterally extensive (~2,000 km
2
) dendritic and stromatolitic carbonate microbialite facies that
extends from ~25 km north of Bristol southward through Taunton (Somerset) and into Seaton
(Devon) (Fig. 1) (Owen, 1754; Hamilton, 1961; Mayall and Wright, 1981; Wright and Mayall,
1981). Alternating episodes of dendrolite and stromatolite are arranged into lenticular,
biohermal mounds (~20 cm thick, decimeters to meters in diameter) separated by a mudstone to
wackestone matrix. Its mammilated surface and internal tree-like architecture have made the
microbialites—known regionally as the ‘Cotham Marble’, although it is a limestone—a geologic
curiosity around the SW UK since at least the 18
th
century (Owen, 1754); polished slabs are
commonly termed “landscape stone” (e.g., Short, 1903). The most recent review of the origins
of the Cotham Marble suggests it resulted from the interplay of sedimentation, bioturbation, and
commensalism between an algal community and serpulid worms in a coastal lagoonal setting
(Wright and Mayall, 1981).
Determining the morphogenesis of the Cotham Marble is of particular interest given its
unique mesomorphology and occurrence within the Triassic-Jurassic ‘boundary interval’
(Hesselbo et al., 2004; Mander et al., 2008), a time of global environmental and biotic crises
(Raup and Sepkoski, 1982; Benton, 1993; Hesselbo et al., 2002; Hautmann, 2004; Kiessling et
al., 2007; Wignall et al., 2007; Greene et al., 2012a). Despite the attention the Cotham Marble
microbialites have received (Hamilton, 1961; Wright and Mayall, 1981), the microstructure,
which is necessary to resolve specific depositional environmental and mechanisms of formation,
! 81
is rarely figured. Here, we analyze the meso- and microfabrics of the Cotham Marble and assess
how its microscopic features contribute to and result in the construction of its meso- and
macrostructure. We also discuss the significance of the Cotham Marble with respect to its
expansive areal extent and its occurrence within the Triassic-Jurassic boundary interval.
LITHOLOGY AND ENVIRONMENTAL SETTING
Bioherms of the Cotham Marble are restricted to the top of the Upper Cotham Member of
the Lilstock Formation. The Cotham Member (divided into the Lower and Upper Cotham) is a
~2.5 m thick sequence of sedimentary rocks that succeed the dark laminated shales of the
Westbury Formation (Swift and Martill, 1999). The Lower Cotham Member is composed of thin
interbedded siltstones and fine-grained calcareous sandstone, which is commonly wave-rippled
(Mayall, 1983). A laterally extensive ‘deformed bed’ exhibiting apparent soft sediment
deformation extends for much of the SW UK (Mayall, 1983), and has been interpreted as the
result of seismic activity associated with the opening of the Atlantic Ocean (Swift and Martill,
1999), or deformation caused by an extraterrestrial impact (Simms, 2003; Simms, 2007). The
‘deformed bed’, which marks the boundary between the Lower and Upper Cotham Members, is
capped by a fissured surface with desiccation cracks up to ~90 cm deep (Hesselbo et al., 2004).
The first occurrence of the Jurassic ammonite Psiloceras planorbis in the basal part of the
Blue Lias Formation was previously proposed as the Global Stratotype Section and Point (GSSP)
for the base of the Jurassic and is the principal boundary marker (Warrington et al., 2008). The
occurrence of the Cotham Marble in the Upper Cotham Member indicates the microbialites are
Rhaetian in age. The Upper Cotham Member was deposited during a relative sea level rise
(Hesselbo, et al., 2004) following the sequence boundary denoted by the fissured surface. The
depositional setting of the Upper Cotham Member is interpreted as a shallow coastal
! 82
environment that alternated between periods of restriction and connection to open marine waters,
on a storm dominated carbonate ramp during a marine transgression (Hesselbo et al., 2004).
Unlike the Lower Cotham Member, the Upper Cotham Member (~0.5 m thick) does not exhibit
synsedimentary deformation. At Manor Farm (Fig. 1), the Upper Cotham Member is about 0.5
m thick and exhibits a 2–3 cm thick layer of ferruginous sand followed by 0.5 m of blue-grey
silty mudstone with starved ripples (Simms, 2003, 2007), indicating deposition in shallow water.
In the northern sections, the Cotham Marble is located at the top of the Upper Cotham Member
and is conformably succeeded by shelly mudstone and limestone of the Jurassic Blue Lias
(Radley and Carpenter, 1998; Simms, 2007). Despite their striking lateral extent across the SW
UK (Fig. 1), the Cotham Marble is absent in the well-studied sections along the Severn Estuary
(e.g., St. Audries Bay and Lavernock Point) (Gallois, 2009).
METHODS AND APPROACH
Here we focus on specimens of the Cotham Marble that contain alternating dendritic and
laminated fabrics. Other forms of the Cotham Marble exist, including some that contain a single
layer of dendrolites and others that are associated with a flat pebble conglomerate, possibly
resulting from current reworking before lithification (Hamilton, 1961).
Samples of Cotham Marble analyzed in this study originate from Manor Farm Quarry,
Lower Woods Nature Reserve, Cotham House in Bristol, and Pinhay Bay in the Dorset coast
(Fig. 1). Decimeter size samples were cut, polished, and scanned on a high-resolution scanner
for mesostructural analyses. Microscopic analyses were carried out via petrographic study of 44
thin sections, 24 acetate peels, and observations using a Scanning Electron Microscope (SEM)
equipped with an Energy Dispersive Spectrometer (EDS).
! 83
Total organic carbon (% TOC) measurements targeting specific microbialite fabrics
(dendrolite and fill) were measured using both an elemental analyzer (EA 1110 CHNS-O) and an
EA coupled to a Picarro Cavity Ring Down Spectrometer (G2121-i). About 30 mg of carbonate
powder (per sample) was wetted with 60 µL of de-ionized water, and acidified overnight in HCl
fumes under a slight vacuum. Samples were then dried overnight and the carbon content was
subsequently measured using the CHNS CO
2
analyzer, in the first case, and using the Picarro
CO
2
analyzer in the later case. Both methodologies were standardized and the relative precision
of the TOC measurements, based on control analysis of a lab reference standard for both runs
(Sulfanilamide and L-Glutamic Acid USGS40), was better than 2.5% for the EA run and better
than 2% for the Picarro run. Replicate precision on sample runs is better than 5% for both
instruments.
RESULTS
Mesostructure of Vertical Cross Sections
The Cotham Marble contains two distinctive mesostructural elements—a laminated
mesostructure (L) and a dendritic mesostructure (D)—that alternate with one another to form at
least five growth phases (Fig. 2; L1, D1, L2, D2, L3; see also Wright and Mayall, 1981). While
the same five growth phases can be seen at all sites sampled, the overall thicknesses of individual
mounds varies (~10–20 cm thick). The following mesoscopic observations apply only to
samples from Manor Farm, Lower Woods, and Bristol as we were not able to collect a large
enough sample from Pinhay Bay for mesoscopic analyses due to landslip cover. For comparison,
Figure 4D in Gallois (2007) is a specimen from the Dorset coast containing similar growth
phases.
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Some samples of the first laminated phase (L1, Fig. 2) contain relatively flat laminae to
small-scale cross-bedding (Fig. 3C). Other samples are composed of stromatolitic domes and
cones (Fig. 3), where black, conical laminae alternate with intervals of evenly laminated, cream
to grey colored bands. In some instances, the lighter-colored bands conform to and inherit the
cone shape of the preceding black laminae (Fig. 3A). Overall, there is little evidence of
truncation and the transition into the first dendritic portion of the deposit (D1) appears texturally
gradual.
Layer D1 is initiated with the occurrence of ~2 mm-thick domed microdigitate structures
(“hedges” in Owen, 1754; Hamilton, 1961) that are darker than the preceding laminae of L1.
Their thickness and morphology is extremely consistent across samples from the three sites,
despite the fact that the sites are located over 10 km apart (Fig. 1). Some of the microdigitate
structures give rise to larger dendritic forms (1–5 cm thick) that comprise the framework of the
first dendritic layer. The dendrolites contain a distinct dark outline that contrasts with the lighter
color of the fill (Fig. 2). In some specimens the fill between the dendrolites appears laminated
(Fig. 2C).
D1 is followed by a texturally gradual transition into L2, another laminated phase (Fig.
2). The laminations are darker in color compared to the laminations in the fill of D2. Some
laminations form domes as they drape the preceding tallest dendrolites indicating some
dendrolites had a ~2 cm relief above the seafloor. In some cases the domes result in cones
reminiscent of Conophyton (cf. Mayall and Wright, 1981) with a mesoscopically visible axial
zone (see arrow in Fig. 2A, L2). Additionally, all of the specimens we investigated contain an
undulose lamination within L2 that is darker in color compared to its surrounding laminae,
! 85
exhibiting ~1 mm thick microdigitate structures (see arrows in L2, Fig. 2A–D). The
microdigitate structures resemble the microdigitate structures of D1.
The second dendritic layer, D2, lacks the basal ‘hedges’ observed in D1. Transition into
D2 is texturally abrupt and the developing dendrolites are not well constrained into dendritic
forms but display a rather more thrombolitic texture. The fill contains sand-size detrital grains in
the basal portion of D2 that diminish in quantity towards the top. Like D1, the fill is lighter in
color compared to the dendrolites and some specimens exhibit concave-up laminations (Fig. 2).
Transition into the final laminated phase, L3, is texturally gradual and laminations drape
over the dendrolites. The thickness of the final laminated phase varies from 0.5 cm to about 2
cm.
Mesostructure of Horizontal Cross-Sections
The surface of the Cotham Marble is mammilated (Fig. 4). In some specimens,
uniformly spaced depressions occur between the dendrolites (Fig. 4A; Short, 1903; Wright and
Mayall, 1981). The higher portions correspond to the dendrolites and the depressions correspond
to the fill described in D1 and D2 above. A polished horizontal cross-section through a dendritic
phase demonstrates the dendrolites are spatially arranged into a reticulate network (Fig. 4B).
Microstructure
Laminated Phase 1
Microscopically L1 primarily consists of laminated micrite with some microspar (Fig.
3B). Some regions develop a fenestral and mesh-like fabric that coalesces into cone-shaped
structures such that succeeding laminae result in cones (Figs. 3A–B). The spaces between the
micritic mesh are filled with microspar. The central ‘axial zones’ of the cones typically contain
round to subrounded spar-filled fenestrae (Fig. 3B). Samples that are cross-bedded (Fig. 3C)
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contain incursions of coarser sediment containing larger grains (~300 µm in diameter) including
well-preserved echinoderm fragments (Fig. 3D). The cross-bedded samples also contain rare (~1
mm-thick) calcite pseudomorphs after gypsum displaying a lenticular habit (Fig. 3E). Gypsum
pseudomorphs have been detected in samples from Manor Farm and Bristol and are restricted to
a single layer in L1.
Dendritic Phase 1
The transition into the dendritic phase is gradual as laminations from L1 rise to
microdigitate structures (Figs. 5A and 5E). Most of the small domes are succeeded by a
prominent micritic mesh-like fabric (~1–2 mm thick) (Figs. 5E–F). Space between the micritic
mesh is filled with microspar and spar (Fig. 5F). Following the mesh fabric, micritic laminations
develop once again (Fig. 5E). Most microdigitate structures are approximately 3 mm in height;
however, a few continue vertical laminated accretion, and develop into the conspicuous
mesoscopic dendrolites (Fig. 5A). The internal structure of the dendrolites is dominantly
laminated micrite, but peloidal micritic fabrics also occur (Figs. 5C–D). A semantic argument
could arise here, because what is clearly dendrolitic at the mesoscale, can in many instances be
laminated (e.g., stromatolitic) at the microscale; regardless, we consider these structures
“dendrolites” while acknowledging that some are laminated at the microscale.
Internally, dendrolites contain round to subrounded spar-filled fenestrae (~100 µm–300
µm in diameter) with a prominent dark outline (Fig. 5B). The dark outline is similar in color and
thickness to the dark outline that also envelops the mesoscopic dendrolites (Figure 2; Figure 5A-
B). The internal micritic fabric of the dendrolites contains abundant, largely evenly distributed,
dark, round to rod-shaped, and filamentous-shaped structures (Fig. 6). The dark microstructures
are confined to micritic portions of the dendrolite and do not occur in the microsparitic regions
! 87
(Fig. 6C). SEM and EDS analyses of the dark structures reveal a dominance of Fe and S,
whereas areas that do not contain the structures are dominated by Ca (Fig. 7). The area between
the dendrolites is dominated by microspar and spar. Some regions of the fill contain distinct
clusters of organic-walled cells (Figs. 5F–G; Fig. 8). Clusters usually range from five to twenty
round cells that are on the order of 20–100 µm in diameter (Fig. 8). Analysis of acetate peels
revealed the cell wall is radially-striated (Figs. 8D–E), and some cells contain linear diagonal
sutures (Fig. 8F). Overall, very few clastic grains occur in the fill.
Laminated Phase 2
The first dendritic phase is succeeded by spar and micrite laminations (L2). Laminations
consist of micritic shrubs and the inter-shrub fill is composed of microspar (Fig. 9). A
conspicuous portion of the laminations is an extremely laterally continuous set of laminae that
rise into microdigitate stromatolites (Fig. 2, arrows; Figs. 9A–B). The stromatolites are on the
order of about 1 mm in height and their edges contain a dark outline similar to the outline that
envelops the dendrolites. The space between the stromatolites varies in composition from
micritic shrubs to microspar along the edges of the stromatolite, but most commonly, the space is
filled with microspar and spar. The space between the columns is similar in size and shape to the
branching portions that develop on the ends of the dendrolites (compare Figs. 9B–C).
Dendritic Phase 2
Transition into the second dendritic phase, D2, is abrupt and denoted largely by the
presence of sand-size fossil fragments in a matrix of micrite (Fig. 10). Unlike D1, D2 does not
contain the micritic mesh network (e.g., Fig. 5E). Larger, irregularly shaped fenestrae, like those
present in D1, also occur (Fig. 10B). The fill portion between the dendritic regions contains
many detrital grains including an assemblage of largely intact bivalve shells, fish bones,
! 88
gastropods, and the round organic-walled structures found in D1 (though fewer in number) (Figs.
10B–C). The fill between the dendrolites is laminated near the top and exhibits concave-down
laminations, composed of micrite. The dendrolites are narrower near the top compared to the
base and contain a laminated and peloidal internal texture.
Laminated Phase 3
The top of the dendrolites of D2 are capped by draping dense laminations of micrite, L3.
The draping nature of the laminae suggests a ~2 cm synoptic relief for the dendrolites and
laminations. Some portions of the laminations contain a shrub-like micromorphology as in L2.
Horizontal laminations succeed the drapes and some hand samples display microdigitate
stromatolites (Fig. 2C).
Microstructure of Cotham Marble from Pinhay Bay
Figure 11 illustrates the microstructure of a sample from Pinhay Bay where the transition
from L1 to D1 is preceded by round tubular structures. The tubes can be up to about 500 µm in
diameter (Fig. 11). Tubes can be distinguished from voids that occur within the dendritic
structures based on the fact that the tubes contain a coating of nearly consistent thickness
whereas the round structures within the dendrites contain a dark, diffuse outline.
Horizontal cross-sections
Cross-sections parallel to bedding across D1 display sharp boundaries between the
dendrolites and the fill. Internally, the dendrolites are very similar in structure to the dendrolites
in vertical cross-sections, containing both the round voids and a black outline. Areas that contain
spar and microspar are strictly in the fill portions between the branching dendrolite as well as
filling-in the rounded fenestrae (Fig. 12). The interconnected nature of the dendrolites is
apparent, delineating the polygonal shapes of the fill (Fig. 12A). Horizontal cross-sections also
! 89
expose a complex branching pattern within the dendrolites where spaces between adjacent
branching regions maintain a relatively consistent diameter of about ~200 µm (Fig. 12B).
Total Organic Carbon
The TOC results on the fill and dendrolite fabrics are listed in Table 1 and plotted in
Figure 13. The mean dendrolite TOC is 0.22% (n=14) and the mean fill TOC is 0.17% (n=10).
The average TOC value for the dendrolites is higher than that for the fill (t=2.66, p<0.05).
DISCUSSION
Morphogenesis: Previous Interpretations
Previous studies suggest the dendritic layers of the Cotham Marble resulted from a
commensal relationship between the problematic worm-like microfossil Microtubus communis
and algal microorganisms. The observation of round ellipsoidal voids in the dendritic structures
led previous authors to suggest the Microtubus-algal association gave rise to the dendritic growth
forms (Wright and Mayall, 1981). However, here we suggest the construction of the spar-filled
voids within the dendrolites is better explained via alternative mechanisms. We found
Microtubus to be widespread and conspicuous only in our samples from Pinhay Bay, thus
bringing into question its previously proposed integral role in the formation of the dendritic
fabric. Furthermore it seems unlikely that the presence of serpulid worms would be some of the
primary agents in constructing the dendrolites, especially in light of the consistent and laterally
continuous dendritic and laminated phases (Figs. 1 & 2).
Additionally previous work invokes the presence of burrows in the laminations of the
second and third laminated phase (Wright and Mayall, 1981). The burrowed hypothesis suggests
the microbialite was present and then burrowed, where the burrows create a network of spaces
between once continuous microbialite. However, microbialite in between the putative burrows
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are best described as a series of branching microdigitate stromatolites, each revealing convex
lamination$and lacking truncation expected during burrowing (Fig. 9). Thus, the present
morphogenesis hypothesis is at odds with the observations we have presented. Below is a
description of an alternative interpretation of the growth history of the Cotham Marble, starting
with L1.
Morphogenesis: New Interpretations
Samples investigated in this study, indicate two dominant modes of deposition for L1: (1)
mesoscopically cross-bedded laminae (Fig. 3C); and (2) cone-forming laminae (Figs. 3A–B).
The sharp contacts in the cross-bedded laminae indicate very localized regions that were
susceptible to episodes of desiccation (Fig. 3E, the presence of rare gypsum pseudomorphs) as
well as input of relatively coarser sediment (Fig. 3D) possibly resulting from low energy currents
or storm activity. In samples from Manor Farm, the presence of well-preserved echinoid
fragments just a few centimeters below the start of the first set of dendrolites (Fig. 3D) is notable
as echinoderms are stenohaline organisms requiring normal salinity marine waters (Stickle and
Diehl, 1987). Although there is a possibility that the fragments were transported into the
microbialite depositional environment during a storm event, the lack of evidence for extensive
transport (given the pristine nature of the fragments) indicates that open marine waters were
nearby. Additionally, echinoid fragments have been found just centimeters below the occurrence
of the Cotham Marble at other sites outside of Bristol (Duffin, 1980), further supporting the
proximity or occurrence within normal marine conditions. However, the presence of one thin
layer of evaporite casts (Fig. 3E) within these cross-bedded laminae suggests a brief period of
deposition in a hydrologically closed setting before the onset of the dendritic phases.
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Cone-forming laminae are composed of mesh-like and shrub-like micritic layers where
the space between the shrubs and mesh is filled with isopachous spar and microspar indicating
the spar cement formed within open space in the mesh early, before compaction. The mesh and
shrubs coalesce to form cone structures that resemble Conophyton (e.g., Wright and Mayall,
1981; Fig. 3). Succeeding laminations rise to a point and display significant inheritance (Fig.
3A). Similarly-shaped fenestrae associated with axial zones of cone-shaped stromatolites have
been interpreted as relicts of oxygen bubbles produced by photosynthetic bacteria (Bosak et al.,
2009; Bosak, et al., 2010). Photosynthetic mats help stabilize the bubbles and eventually
cements fill the bubble space (Bosak et al., 2010). However, for the bubble morphology to be
preserved, cementation must be extremely rapid, before the mat has time to decay, suggesting the
microbialites were forming in an environment that fostered rapid carbonate precipitation.
Despite differences in mesofabrics between the two types of L1 laminations, both exhibit
an overall dominance of micrite, indicating deposition in a quiet water, low energy regime. The
microbial fabrics of the cone-forming laminae however, contain a higher degree of in situ
precipitation (perhaps associated with EPS calcification, cf. Riding, 2000), whereas the cross-
bedded laminae contain a greater component of trapped and bound micrite. L1 therefore formed
in very quiet water where some localized regions were dominated by in situ microbial carbonate
and others contain a greater component of pelagic sedimentation.
Dendrolites
The dendritic portions are preceded by an extremely laterally extensive (>10 km), but
gradual shift from the relatively flat laminations of L1 to dome-like vertical growths of
microdigitate columns from which a micritic mesh develops (Figs. 5C–D). These micritic
filament-like features have been the subject of much speculation and are one of the dominant
! 92
reasons that led to the first biologic hypothesis for the origin of the Cotham Marble (Hamilton,
1961; Mayall and Wright, 1981; Wright and Mayall, 1981). Their resemblance to tuft structures
(Horodyski, 1977) led Wright and Mayall (1981) to interpret the ‘hedges’ as algal tufts that result
from the clumping activity of motile algae displaying phototaxis (e.g., Walter et al., 1976).
Similar structures have been observed on the surfaces of oxygen rich bubbles within laboratory
photosynthetic mats, resulting from the production of oxygen bubbles within the mat (Fig. 5–6;
Bosak et al., 2010). Mata et al., (2012) suggest the production of gas within microbial mats
could cause the mat to rise and foster orientation of filament bundles. Thus, phototaxis is not the
only mechanism that can lead to filament orientation. Furthermore, the extreme lateral
continuity of these ‘tuft’ structures across southwestern England suggests that, despite their
likely microbial origin, their formation is significantly influenced by a widespread environmental
change following deposition of the preceding relatively flat laminae of L1.
The dendrolite-fill boundary is sharp, delineated by a black outline on the dendrolites and
a change in grain size (Fig. 5B). We interpret the dark, conspicuous round to filamentous
structures as putative microfossils that could be the primary agents in the construction of the
dendrolites based on: (1) their consistent size; (2) even distribution within the dendrolites; and
(3) restriction to the dendritic regions (Fig. 6). Our SEM EDS analyses of the putative dendritic
microfossils indicate they are coated in iron and sulfur (likely pyrite) (Fig. 7). The pyrite
possibly resulted from the anaerobic degradation of the bacterial organic matter possibly by
sulfate reducing bacteria (e.g., Canfield, 1989; Schieber, 2002). Anaerobic pyritization occurred
after the mat was buried, but early enough to preserve the rod-shaped morphology of the
microorganisms. The prominent black round structures that occur in association with the
filaments and rods (Figs. 6–7), could be micrometer scale porosity structures associated with the
! 93
presence of bacteria (e.g., Bosak et al., 2004) and/or the mineralized byproducts of the bacterial
decay of organic tissues (e.g., Schieber and Arnott, 2003). The morphological preservation of
the microfossils is possibly the result of early pyritization and subsequent rapid calcification in
waters with anomalously high saturation with respect to calcium carbonate.
The round spar and microspar-filled structures like those in Figure 5B have been
previously interpreted as the microfossil of Microtubus communis, suggesting the vertical nature
of the dendrolites is controlled by the orientation of Microtubus (Wright and Mayall, 1981).
However, we found Microtubus to be a conspicuous component of the dendritic phases only in
the samples from Pinhay Bay where tubular structures occur outside the dendrites in the fill (Fig.
11). Microtubus is a marker fossil indicating central parts of the Norian and Rhaetian Tethyan
reefs (Flügel, 2010), thus corroborating its occurrence within these microbialites. However, our
study suggests that the interpretation of Microtubus as one of the primary mechanisms of
formation of the dendrolites may be problematic based on the fact that (1) Microtubus tubes as
defined here do not occur within the dendrolites and therefore cannot be assisting in the vertical
growth of the dendrolites and (2) we found Microtubus to be a significant component of the
microbialites only in our samples from Pinhay Bay (Fig. 11).
To clarify morphologic similarities between the various spar and microspar-filled
structures that occur in the dendritic phases we present images of each of the different structures
at the same scale in Figure 14. Although the structures resemble one another in shape and size,
we describe at least three possible origins. Microtubus communis (Fig. 14A) occurs in greatest
concentration in the samples from Pinhay Bay (Fig. 11). Microtubus can be distinguished from
other spar-filled structures based on the presence of a micritc rim of consistent thickness and its
occurrence primarily in the fill. We found two distinct types of structures that occur within the
! 94
dendrolite matrix. Round to elliptical structures that contain a dark rim are likely the result of
evenly-spaced, complex three-dimensional branching within the mats that results in round to
elliptical voids when represented two-dimensions (Figs. 5B and Fig. 12 and Fig. 14B). A third
type of spar and microspar-filled structure occurs within the microbialite matrix but does not
contain a dark outline (Fig. 14C). The similarity in shape, size, and strict association with the
microbialite matrix suggests they can possibly result from the production of gas bubbles (sensu
Bosak, et al., 2010; Mata, et al., 2012) potentially resulting from metabolic processes associated
with the putative microfossils (Fig. 6). Similar round to elliptical structures form within cone-
forming photosynthetic bacterial mats (Bosak et al., 2010). The modern mats produce round
voids whose outline represents the microbial community and corresponding organic matrix that
was displaced and condensed during the production of the bubble (Bosak et al., 2010). Round to
ovoid shaped structures that are now filled with spar (and suggesting they were once open space)
are present in the internal structure of many of the dendrolites we have investigated and
strikingly resemble those found in L1 (Fig. 3B; Fig. 14D) suggesting they formed via similar
mechanisms. The gas bubble and microbial mat branching types of structures can be difficult to
differentiate from one another due to their mutual occurrence within the microbialite matrix (e.g.,
Fig. 14C), but the careful use of repetitive acetate peel analyses assists in highlight the true
length of the structures.
Tube Structures
Following the microdigitate domes, the propensity for the laminations to grow vertically
rather than horizontally is further propagated giving rise to the mesoscopic dendrolites (Fig. 2).
Horizontal cross-sections reveal the dendrolites are spatially arranged on their respective mounds
into a macroscospic mesh-like fabric where, while being interconnected, adjacent dendritic
! 95
fabrics are separated from one another by ~1–3 cm round to polygonal spaces, resembling a
honeycomb (Owens, 1754; Short, 1903; Wright and Mayall, 1981)(Fig. 4). These bedding plane
views of the microbialite highlight their close resemblance to tube structures (Corsetti and
Grotzinger, 2005). A sample of the surface of the Cotham Marble demonstrates the dendrolites
had relief of up to ~2 cm and they appear to have been accreting faster than the intra-dendrolite
fill, giving the impression of holes at the dendrolite-water interface when they were forming
(Fig. 4A).
Tube structures represent growth in quiet water resulting from in situ precipitation of
carbonate in environments that are supersaturated with respect to calcium carbonate (Corsetti and
Grotzinger, 2005 ). Rather than forming discrete tree-like structures (as it appears in cross-
sections perpendicular to bedding), they form an interconnected network of dendritic
microbialite whose associated ‘tubes’ are subsequently filled often resulting in concave up
laminae (Fig. 2C). The isopachous cement in the space between the dendrolite branches suggests
the space remained open as succeeding layers of dendrolite developed, but the cement formed
before the dendrolites were compacted. The conspicuous contrast in color and grain size
between the dendrolite and the fill as well as their significantly different TOC content (Fig. 13)
indicates their distinct origin, where the darker and finer grained dendrolites likely reflect a
greater organic component (Fig. 6).
While reticulate and tufted mats can be found in modern environments (e.g., Walter, et
al., 1976; Wharton, 1994), and the reticulate framework can be manipulated in laboratory
cultures (Bosak, et al., 2010; Shepard and Sumner, 2010), the preservation of such structures via
rapid lithification is not readily observed in the modern. Thus, the discovery of well-preserved
tube structures in the rock record is likely an indication of high saturation state of calcium
! 96
carbonate needed to lithify the intricate shrub, reticulate, and dendritic fabrics quickly (e.g.,
Fraiser and Corsetti, 2003; Shapiro and Awramik, 2006; Flannery and Walter, 2011). The even
spacing and polygonal intra-host rock shape is probably best explained by space limitation
associated with competition for limited resources (Petroff et al., 2010). Furthermore, the
organization of reticulate mats (e.g., Bosak, et al., 2010; Shepard and Sumner, 2010) might be a
prerequisite to the development of tubestones given the remarkable resemblance of modern
reticulate mats to plan view polished sections of tubestones (Fig. 4).
‘Burrows’
Laminated phase L1 and L2, contain microdigitate structures that have been previously
interpreted as the burrowing activity of metazoans (Wright and Mayall, 1981). Here we suggest
the mm-scale depressions are better explained by the formation of microdigitate stromatolites
(see references in Riding, 2011) and therefore reflect a growth process as opposed to a
destructive process. Comparing the mm-scale depressions to the dendrolite branches (Fig. 9)
suggests they formed in a similar manner. While the structures indeed resemble burrows in
mesoscopic images (Wright and Mayall, 1981), in thin section, photomicrographs reveal they are
best explained by a microbial microdigitate growth based on: (1) their resemblance in size,
shape, and crystal size/texture to branching portions of the dendrolites; and (2) their tendency to
rise from the same lamination as opposed to forming depressions that vary in depth.
Additionally, the microdigitate structures are very laterally continuous (Fig. 2 arrows), being
present in the samples we investigated, ~25 km apart, suggesting a significant environmental
control on the structures as opposed to resulting from a localized phenomenon.
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Scale of Control
One of the most prominent features of the Cotham Marble microbialites is their
consistent mesomorphology across tens of kilometers (Woodward, 1892). The representation of
the same growth episodes across tens of kilometers suggests the same conditions that affected the
growth and development of either dendritic and or laminated fabrics were the same across a vast
distance (at least 2,000 km
2
). The occurrence of the same mesostructure from site to site implies
a well-mixed depositional environment that extended for tens of kilometers, giving rise to the
same overall morphology. Such conditions are possible given dominance of in situ chemical
precipitation, as opposed to formation via trapping and binding of detrital grains. The similarity
in mesostructure also implies that large-scale processes such as water chemistry and/or climate
played a fundamental role in controlling the development of the alternating growth phases (Fig.
2).
A high degree of similarity across large distances indicates a greater likelihood of large-
scale influences dominantly controlling the fabrics. The laminations for example, exhibit a
greater degree of large-scale control given how similar they are in thickness and number of
laminations. The dendrolites, on the other hand, vary in size, structure, and degree of branching
from dendrolite to dendrolite at distances of only one or two cm emphasizing the relatively
greater importance of cm to mm scale influences, which is supported here by the presence of
microfossils. Nonetheless, the dendrolites are present over a vast areal extent indicating that
similar biological communities were synchronously dominant across the region. While several
influences at different scales contribute to the resulting microbialite micromorphology (Burne
and Moore, 1993; Riding, 2000; Dupraz et al., 2009), multiscale analyses performed in
! 98
conjunction with observations of lateral extent allow us to better assess the dominant scales of
control (local versus nonlocal) on the fabrics.
Tasmanites
We interpret the round, organic-walled structures that occur in the fill of D1 and D2 (Fig.
8) as prasinophyte algal cysts of the genus Tasmanites. Their range in size from (~20 to 100
µm), prominent thick wall, radially arranged pore canals within the wall (Figs. 8D–E), and the
linear suture (Fig. 8) are characteristics of Tasmanites cysts (Wall, 1965; Parke et al., 1978;
Tappan, 1980). Tasmanites is an algal prasinophyte with a motile quadriflagellate cell phase and
a cyst phase (phycoma) in which new motile cells develop (Tappan, 1980). Prasinophyte algae
have been termed ‘disaster species’ for their abundance during episodes of biotic crises (Tappan,
1980; Schwark and Empt, 2006; Schootbrugge et al., 2007; Jia et al., 2012), possibly due to their
ability to return to their phycomata phase when conditions are unfavorable (Tappan, 1980).
Prasinophyte peaks have been recorded in other Tethyan Triassic-Jurassic sections (Kuerschner
et al., 2007; Götz et al., 2009; Bonis et al., 2010) during the end-Triassic extinction interval. The
occurrence of abundant prasinophytes within the Cotham Marble may be related to the peak in
prasinophytes seen in other Tethyan sections during the End-Triassic extinction interval or they
may reflect deposition in a restricted setting.
A unique aspect of the detection of Tasmanites in this study is their occurrence in
clusters. Most palynomorph studies report prasinophytes from maceration or biomarker
analyses, which result in loss of meso- to micro- depositional context. Here, their detection in
thin section and acetate peel analyses reveal their life position before burial allowing us to
ascertain their restriction to the fill (and not within the dendrolites) and tendency to occur in
clusters (Figs. 8A–D). Hamilton (1961) describes the presence of ‘bulbous filaments’ near the
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‘hedges’ where we detect the highest concentration of the Tasmanites clusters, and invokes these
‘bulbous filaments’ as contributors to the growth of the hedges and dendrolites. However, it is
clear from our analyses that although Tasmanites dominantly occur in the bottom portions of D1,
Tasmanites occur in the fill, and rarely within the dendrolites.
Relevance to the End-Triassic Mass Extinction
The end-Triassic mass extinction is one of the ‘big five’ mass extinctions of the
Phanerozoic (Raup and Sepkoski, 1982), which has been temporally linked to the Central
Atlantic Magmatic Province (CAMP) (Marzoli et al., 2004; Schoene et al., 2010; Blackburn et
al., 2013). The release of CO
2
and other volatile gases associated with the emplacement of
CAMP—the most aerially extensive Large Igneous Province of the Phanerozoic (McHone,
2003)—coincides with a negative carbon isotope excursion used as a global stratigraphic marker
for the Triassic-Jurassic boundary (Pálfy et al., 2001; Ward et al., 2001; Guex et al., 2004; Črne
et al., 2011).
A series of cascading effects likely resulted from the emplacement of a large igneous
province (LIP) like CAMP (Greene et al., 2012b). Orbitally tuned pCO
2
records interbedded
with CAMP flood basalts suggest that CAMP emplacement occurred in large pulses (Schaller et
al., 2011; Schaller et al., 2012), releasing huge quantities of CO
2
, and perhaps induced temporary
ocean acidification (Greene et al., 2012b; Hönisch et al., 2012). However, this effect would have
been temporary, as weathering (enhanced by CO
2
-induced warming) delivers alkalinity to the
ocean, which would promote a supersaturation overshoot with respect to calcium carbonate
(Kump et al., 2009; Greene et al., 2012b; Hönisch et al., 2012), an explanation that has already
been invoked to help explain anomalous early diagenetic carbonate diagenesis at the St. Audries
Bay Triassic-Jurassic transition (Greene et al., 2012a). A sustained warm period together with
! 100
oversaturation of seawater with respect to calcium carbonate would provide ideal conditions for
the development and rapid lithification of microbialites.
Furthermore, the occurrence of the Cotham Marble microbialites at the same level as
what has been interpreted as a post-extinction ‘dead zone’ based on paleoecological studies on
rocks from correlative sections in St. Audries Bay and Lavernock Point (Fig. 1; Mander et al.,
2008), suggests grazing activity by metazoans would have been reduced. Thus, as has been
observed in response to other Phanerozoic biotic crises (Kershaw et al., 2011; Mata and Bottjer,
2011) we envision that the Cotham Marble represents a shift from skeletal carbonate to microbial
carbonate.
Stratigraphically, the Cotham Marble microbialites occur at the top of the Cotham
Member beds $at the level of the initial negative carbon isotope excursion (Hesselbo, et al.,
2002; Hesselbo et al., 2004) that is used to correlate the placement of the Triassic-Jurassic
boundary with other sections around the world. Therefore, the Cotham Marble formed during a
time of significant global environmental and biologic perturbations and could represent a
‘disaster form’ (Schubert and Bottjer, 1992) of the end-Triassic mass extinction (Radley et al.,
2008). We suggest that the microbialites’ striking lateral extent (Fig. 1) and the widespread
occurrence of the prasinophyte Tasmanites phycoma$ interestingly, also dominant at the level
of the initial negative CIE from other Upper Triassic Tethyan sections (Kuerschner et al., 2007;
Götz et al., 2009; Bonis et al., 2010)$ indicates the microbialites capture extensive and
significant petrographic and geochemical information that link them to the end-Triassic
extinction.
! 101
CONCLUSIONS
Textural and petrographic analyses of Upper Triassic carbonate microbialites from the
southwestern United Kingdom reveal evidence of microbial fabrics that resulted from
precipitation of carbonate in waters with anomalously high carbonate saturation. Round to
ovoid-shaped, spar-filled structures within the dendrolites are the product of complex three-
dimensional microbial mat branching and/or metabolic gas production within the mats. The
dendritic texture is a product of rapidly accreting and rapidly lithifying reticulate mats that led to
the formation of the tubestone morphology. The consistent mesomorphology across tens of
kilometers suggests the waters in which the microbialites developed lacked the presence of large
metazoan grazers—whether because of local restriction, or, more likely, the environmental
stressors associated with the end Triassic extinction—allowing for largely uninterrupted growth.
Additionally, the expansive lateral continuity of similar textural phases indicates a dominance of
large-scale controls (e.g., climate, water chemistry, etc.) on microbialite formation. The
occurrence of the microbialite mounds at the level of the Triassic-Jurassic interval along with
their associated Tasmanites acme indicate the Cotham Marble microbialites capture some of the
unusual environmental conditions associated with the end-Triassic of the SW UK.
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FIGURES
Figure 1. Map of the southwestern United Kingdom showing the locations of the Cotham Marble
microbialites. Numbered locations contain examples examined in this study. Locations a-c
(Richardson, 1911); Locations d and h (Hamilton, 1961); Location e (Hesselbo et al., 2004);
Location f (Duffin, 1980); Location g (Vaughan, 1903). Map after Mander et al., 2008. Grey is
Triassic-Jurassic strata.
! 112
Figure 2. The five most common phases of the Cotham Marble labeled L (Laminated) and D
(Dendritic). Specimens A-C are from localities near the Cotham House in Bristol and specimen
D is from Manor Farm Quarry near Aust. Arrows indicate potential coeval laminae.
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Figure 3. Scans and photomicrographs of L1. (A) High-resolution scan of dark, cone-forming
laminae (arrow). (B) Photomicrograph of cone-shaped structures (bottom) and overlying
succeeding laminations. (C) Scan of laminations with small-scale cross-bedding. (D) Two well-
preserved echinoid fragments (arrows). (E) Calcite pseudomorphs after lenticular gypsum
showing dendritic putative microbial encrustations. Samples A–C are from Bristol; Samples D–
E are from Manor Farm.
1cm
1cm
1mm
C
B A
D E
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Figure 4. Bedding plane view of the Cotham Marble. (A) Sample indicates the dendrolite fabric
and adjacent hollow fill; Bristol Museum and Art Gallery, number Cb 4130. (B) Polished
horizontal cross-section through D1 showing the interconnected network of dendrolite separated
by polygonal-shaped fill. Sample A is from Bristol and Sample B is from Lower Woods.
1cm
1cm
A B
! 115
Figure 5. Scan and photomicrographs of D1. (A) Dendrolite mesostructure and adjacent fill. (B)
Photomicrograph of dendrolite and adjacent fill. (C) Dendrolite with laminated internal structure.
(D) Dendrolite with peloidal internal microscopic structure. (E) L1-D1 boundary (F) Micritic
‘mesh’ overlain by micrite laminations. (G) Higher magnification of fill portion of (F) showing
cluster of putative Tasmanites phycomata. Samples A–D are from Bristol and E–G are from
Manor Farm.
1cm 1mm
Dendrolite
Dendrolite
G
micritic,‘mesh’
C B A
D
E
F G
2,mm
! 116
Figure 6. Putative microfossils in the dendrolites. (A) Photomicrograph of a dendrolite branch.
(B) Close-up of dendrolite branch surrounded by fill. (C) Filamentous structures within
dendrolite branch. Sample location: Lower Woods.
Figure 7. Energy dispersive spectrometry analyses of the filamentous microstructures indicating
regions replete with black structures (labeled ‘pyrite-coated filaments’ in Fig. 6C) contain Fe and
S whereas regions that lack black structures are dominated by Ca.
! 117
Figure 8. Putative Tasmanites phycomata. (A) Fill between two dendritic branches containing a
cluster of Tasmanites. (B) Higher magnification of cluster in (A). (C) Photomicrograph of
Tasmanites phycoma cluster. (D) Acetate peel photomicrograph of a cluster of cells illustrating
their radial wall canals. (E) Higher magnification of (D). (F) Photomicrograph illustrating linear
diagonal sutures across putative Tasmanites phycomata. Samples A–C are from Bristol and D–F
are from Manor Farm.
B
linear(suture radial(wall(
canals
C
B A
D
E F
! 118
Figure 9. Laminations of L2. (A) Dark microdigitate structures succeed shrub laminae. (B)
Branching microdigitate structures from the same layer as those in (A). (C) Branching dendritic
structures resembling those in (B). Sample location: Bristol.
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Figure 10. Growth phase D2. (A) High-resolution scan of D2 highlighting detrital grains
between the dendrolites. (B) and (C) illustrate a high concentration of biogenic and detrital
grains in the fill between the dendrolites. Sample A is from Bristol and samples B–C are from
Manor Farm.
1cm
detrital(
grains
A
1mm 1mm
C B
Dendrolite
Dendrolite
! 120
Figure 11. Thin section photomicrograph of a sample from Pinhay Bay. Note the presence of
Microtubus across the bottom of the sample.
Figure 12. Photomicrographs showing complex microbial mat branching. (A) Photomicrograph
of dendritic structure. (B) Higher magnification of (A) illustrating evenly-spaced microbial mat
branching. Sample locality: Manor Farm.
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Figure 13. Box and whisker plot of percent total organic carbon for the dendrolite and fill fabrics.
The whiskers represent the range, the box represents the upper and lower quartiles and the black
horizontal line within the boxes denotes the median. Dendrolite and fill TOC averages are
0.21% and 0.17%, respectively. The standard deviation between duplicate samples was better
than 0.02.
Figure 14. Various spar and microspar-filled ovoid structures. (A) Microtubus from Pinhay Bay.
(B) Complex branching pattern in dendritic matrix from a Manor Farm sample. (C) Ovoid
structures within dendritic matrix in a sample from Manor Farm. (D) Round to irregular-shaped
spar-filled structures within L1 in a sample from Bristol. Scale bar = 500 µm.
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Table 1. TOC measurement results for the dendrolite and fill fabrics. Samples with an asterisk
(*) indicate samples that were analyzed on the elemental analyzer coupled to the Picarro
instrument; all others were measured on the elemental analyzer. The (**) indicate samples
whose measurements were not duplicated.
Fabric Type
Sample ID % TOC
Average of
duplicate
samples
Dendrolite
LWd1 0.165 0.17
LWd1-2 0.174
LWd2 0.191 0.20
LWd2-2 0.204
LWdb 0.220 0.22
LWdb-2 0.220
MFd1 0.149 0.15
MFd1-2 0.146
MFd2** 0.190
MFd3 0.218 0.23
MFd3-2 0.233
MFd4** 0.173
LWD1hedge* 0.239 0.24
LWD1hedge-d* 0.230
LWD2* 0.234 0.23
LWD2-d* 0.217
LWD2hedge* 0.237 0.23
LWD2hedge-d* 0.222
LWDtop* 0.360 0.37
LWDtop-d* 0.381
MFDhedge* 0.233 0.24
MFDhedge-d* 0.237
MFDL2* 0.199 0.21
MFDL2-d* 0.222
BDL2* 0.196 0.20
BDL2-d* 0.202
Fill
MFf1 0.126 0.12
MFf1-2 0.120
MFf2 0.124 0.12
MFf2-2 0.122
LWf1 0.160 0.15
LWf1-2 0.139
LWf2 0.155 0.16
LWf2-2 0.159
LWF1* 0.179 0.18
LWF1-d* 0.188
LWF2* 0.187 0.19
LWF2-d* 0.194
LWF3* 0.199 0.19
! 123
LWF3-d* 0.180
LWF4* 0.205 0.20
LWF4-d* 0.196
LWF5* 0.186 0.191
LWF5-d* 0.196
LWF6* 0.176 0.180
LWF6-d* 0.184
Dendrolite
mean
0.22
1σ
0.05
Fill mean
0.17
1σ
0.03
! 124
Chapter 5:
A widespread microbial carbonate response across the end-
Triassic mass extinction, southwestern United Kingdom
ABSTRACT
The end-Triassic mass extinction is linked to the eruption of the Central Atlantic
Magmatic Province (CAMP), however details about the sedimentary and biological dynamics of
the extinction and subsequent recovery are poorly understood. Here we document the
widespread distribution of carbonate microbialites across the southwestern United Kingdom (SW
UK) that represent a combined environmental and biological response to an ecosystem collapse
across the end-Triassic. The microbialites (1) occur at the same stratigraphic level as the mass
extinction (2) contain a negative carbon isotope excursion in δ
13
C
org
and (3) co-occur with a
bloom of ‘disaster taxa’ prasinophyte algae also dominant in other European sections in
synchrony with the so-called ‘initial carbon isotope excursion’. Further, the microbialites are
composed of intricate microbial mat textures and correlate with mixed carbonate-siliciclastic
facies that contain abundant, well-preserved filamentous microfossils whose unusual
preservation implies growth in waters with anomalously high carbonate saturation. High
saturation state with respect to calcium carbonate may have resulted from a period of enhanced
continental weathering in response to higher pCO
2
levels triggered by the emplacement of the
CAMP. These findings indicate the end-Triassic strata of the SW UK capture a significant
microbial carbonate sedimentary response as is commonly observed across other episodes of
global biotic crisis.
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INTRODUCTION
Microbialites—organosedimentary deposits that have accreted as a result of a benthic
microbial community trapping and binding grains and/or forming the locus of mineral
precipitation (Burne and Moore, 1987)—are significant geobiological indicators of various
environmental and ecological parameters (e.g., seawater chemistry, microbial mat communities,
etc.) (Riding, 2011). Although microbialites occur in modern marine settings, they are
morphologically distinct from their Proterozoic counterparts. Modern microbialites are coarse-
grained and poorly laminated, whereas Proterozoic microbialites are primarily composed of fine-
grained stromatolitic and/or thrombolitic textures (Awramik and Riding, 1988). In the
Phanerozoic, extensive deposits of fine-grained microbialites are most common in
environmentally restricted settings (e.g., alkaline lakes, thermal springs, etc). However, during
times of biotic crisis, fine-grained microbialites expand into normal marine environments
(Schubert and Bottjer, 1992; Mata and Bottjer et al., 2011, and references therein), oftentimes
occurring in the immediate aftermath of ecological disturbance. In these instances, microbialites
can serve to elucidate critical clues about water/atmospheric chemistry, microbial communities,
and sedimentary processes during times of biotic and environmental crises.
The end-Triassic extinction, one of the ‘big five’ mass extinctions of the Phanerozoic
(Raup and Sepkoski, 1982), has been linked to the eruption of the Central Atlantic Magmatic
Province (CAMP) (Blackburn et al., 2013). High pCO
2
levels (McElwain et al., 2009; Schaller
et al., 2011) and the release of volcanic volatiles via massive CAMP eruptions resulted in
widespread warming and the spread of anoxia in shallow marine environments (Richoz et al.,
2012; Jaraula et al., 2013) likely contributing to the collapse of terrestrial and marine
ecosystems. Some of most well-known shallow marine Triassic-Jurassic sections are those of
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the United Kingdom (see Warrington et al., 2008) for which there are detailed stratigraphic
(Hounslow et al., 2004; Ruhl et al., 2010) and biogeochemical records published (Hesselbo et al.,
2004; van de Schootebrugge et al., 2007; Korte et al., 2009; Whiteside et al., 2010; Paris et al.,
2010; Jaraula et al., 2013). Across the southwestern United Kingdom (SW UK), the mass
extinction interval contains a widespread (>2,000 km
2
) carbonate microbialite unit (~20 cm
thick) known regionally as the ‘Cotham Marble’ composed of fine-grained laminated and
dendritic fabrics (Hamilton, 1961), reminiscent of Proterozoic microbialites. The environmental
mechanisms responsible for the widespread development and preservation of the microbialite
unit have been a geologic mystery since the 18
th
century (Owen, 1754). Here we present
combined biogeochemical and petrographic evidence that indicate the microbialites are a product
of at least regional, if not global, environmental controls and contain numerous direct links to the
end-Triassic mass extinction.
STRATIGRAPHY AND ENVIRONMENTAL SETTING
The upper Triassic beds of the SW UK were deposited in shallow epicontinental seaways
connected to the Tethys Ocean to the south (Swift and Martill, 1999). Widespread shallowing
of the depositional system occurred near the top of the Westbury Formation and continued into
the Lower Cotham Member of the Lilstock Formation (Hesselbo et al., 2004). The top of the
Lower Cotham Member contains a widespread fissured horizon (Fig. 1), interpreted as the result
of seismic activity (Hesselbo et al., 2004) or an extraterrestrial impact (Simms, 2007). The
Upper Cotham Member was deposited during a relative sea level rise and contains widespread
ripple marks and mudcracks indicating deposition in very shallow water likely under tidal
influence (Hesselbo et al., 2004). The Cotham Marble microbialites are located at the top of the
Upper Cotham Member but do not occur in the well-studied sections of St. Audrie’s Bay (SAB)
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and Lavernock Point (LP) (Fig. 1). The microbialites, however, are remarkably laterally
extensive (Fig. 2) and contain a widespread distinctive cyclic mesomorphology despite occurring
as discrete m-scale mounds (Fig. 3A). The depositional setting of the Upper Cotham Member
has been interpreted as a normal salinity lagoon (Radley et al., 2008), or a shallow carbonate
subtidal ramp (Hesselbo et al., 2004).
METHODS
Rock samples were collected at high-resolution from the Lilstock Formation of the SAB
and LP sections as well as at various other Cotham Marble microbialite-bearing localities (Fig.
2A). Several specimens of the microbialites were examined at the Bristol Museum and Art
Gallery. Hand size samples were cut, polished, and scanned on a high-resolution scanner for
mesoscopic analyses. Microfacies analyses were carried out via thin section light microscopy
and scanning electron microscopy (SEM). Corresponding thin section billets were micro-drilled
at ~cm-scale intervals for high-resolution stable isotope analyses of carbonate carbon and
oxygen. Sample aliquots of carbonate samples (~30 mg) were used for measurements of δ
13
C
org
.
RESULTS
Geochemistry
Carbon isotope analyses of the bulk organic carbon within the microbialite unit from four
different localities reveal values ranging from -25.8‰ to -29.6‰ (Fig. 2). The overall range in
values spans most of the ‘initial’ carbon isotope excursion (I-CIE) measured at SAB and occurs
at the same stratigraphic level (Fig. 1) based on the most likely correlations between
microbialite-bearing sites and St. Audrie’s Bay (Hesselbo et al., 2004). For most studied
microbialites, the most negative δ
13
C
org
values occur in the first dendritic phase of the
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microbialites, a pattern observed in sites around Bristol and further afield in Pinhay Bay (Fig. 2).
No correlation was found between paired inorganic carbon and oxygen isotopic values.
Petrography
Microscopically, the Cotham Marble microbialites display intricate dendritic branching
patterns and well-preserved fine-grained stromatolitic laminae (Fig. 3B). The interstitial fill
between the dendritic microbialite fabric contains distinct clusters of round, organic-walled cells
(~10-100 µm in diameter) that were detected in microbialites from all of the investigated sites
and also occur in clusters in the Upper Cotham Member at SAB (Fig. 3C-D). Common features
of the organic-walled cells are linear diagonal sutures and radially-arranged wall canals.
Although the microbialites do not occur at SAB and LP, perhaps due to a greater
component of clastic input, the shallow rippled facies of the Upper Cotham Member in SAB and
the correlative heterolithic facies of LP contain copious filamentous microfossils (~10 µm in
diameter and ~100-200 µm in length) (Fig. 4A-C). The filaments are aligned parallel to the
foreset laminae of ripple marks and are associated with sedimentary pyrite, quartz, and
carbonate-rich laminations (Fig. 4A-B). Additionally, mm-scale dendritic fabrics reminiscent of
the Cotham Marble dendrolites occur in the oolitic beds at LP suggesting a possible connection
to the microbialite dendritic phases (Fig. 4D).
DISCUSSION
Microbial Facies
The Upper Cotham Member is characterized by a set of unique microbial features. The
Cotham Marble microbialites coincide laterally with deposition of shallow water facies in SAB
and LP, where the presence of flat-topped ripples, mudcracks, and heterolithic facies of
interbedded sandstone and shale reflect a tidally-influenced setting that underwent brief periods
! 129
of subaerial exposure (Hesselbo et al., 2004). Microscopically, well-preserved filamentous
fossils (Fig. 4) could be filamentous algal mats that developed extensively along the shallow
water facies and were effective in providing cohesiveness to the sediment. The widespread
development of microbial mats during what is interpreted as a post-extinction ‘dead zone’
(Mander et al., 2008), indicates that in the absence of extensive bioturbation, the mats were able
to expand laterally and become preserved. The significant component of sedimentary pyrite
suggests the presence of sulfate reducing bacteria below the photosynthetic surface mat layers
(e.g., Schieber, 1998a). Although the taxonomic affinity/metabolic pathways employed by the
filamentous fossils are unknown, their exceptional preservation predominately within carbonate-
rich laminae (Fig. 4A-B) suggests the filaments were calcified quickly inhibiting complete
decay, whether because of high rates of sulfate reduction within the sediments and/or high
seawater carbonate saturation.
The laterally extensive bloom of organic-walled, round microfossils co-occurs with
deposition of the Cotham Marble microbialites and Upper Cotham Member facies at SAB (Fig.
1). Radially arranged wall canals, linear diagonal sutures, and their occurrence in blooms
strongly suggests the fossils represent phycoma clusters assignable to the prasinophyte green
algae Tasmanites, known to have bloomed later in the section at SAB (van de Schootbrugge et
al., 2007) and across the I-CIE excursion in other Triassic-Jurassic sections across Europe (Pálfy
et al., 2007; Götz et al., 2009).
Carbon isotopes
The shift in δ
13
C
org
from about -25.8‰ to -29.6‰ within the microbialites (Figs. 1 and
2C) corresponds in absolute value and magnitude to the I-CIE from SAB and occurs at the same
stratigraphic level (Hesselbo et al., 2004), indicating the onset of the first dendritic phase of the
! 130
Cotham Marble contains what has been interpreted as an I-CIE. Interestingly, the first dendritic
phase includes the greatest abundance of the organic-walled microfossils (Fig. 2B) interpreted as
Tasmanites phycomata (Ibarra et al., 2014b). Our observations are similar to analyses made with
the onset of a second ‘main’ (Hesselbo et al., 2002) CIE at SAB that also coincides with an acme
of Tasmanites (van de Schootbrugge et al., 2007). The occurrence of the I-CIE at the same level
as peaks in green algae in other end-Triassic sections (Götz et al., 2009; Palfy et al., 2007;
Kuerschner et al., 2007; Ruhl et al., 2010) suggests a possible causal relationship, whereby the
CIE in bulk organic matter is the result of organic matter compositional changes (van de
Schootebrugge et al., 2008). However, compound specific carbon isotope analyses from Austria
demonstrate a lack of significant changes in the sedimentary organic matter that would indicate a
causal relationship between a peak in green algae and a synchronous δ
13
C
org
excursion (Ruhl et
al., 2010). While more work is needed to resolve the nature of CIEs associated with pronounced
peaks in green algae across other end-Triassic sections (e.g., Jaraula et al., 2013), green algal
blooms appear to be a widespread phenomenon across European end-Triassic sections
highlighting a strikingly similar response to other episodes of biotic crisis (Jia et al., 2012, and
references therein).
Significance of microbialite and associated biogeochemical observations
Two hypotheses have been invoked for the extensive development of microbialites
during times of biotic crisis: (1) opening of niches previously occupied by organisms affected by
the extinction together with suppression of grazing/bioturbation (e.g., Schubert and Bottjer,
1992) and (2) changes in carbonate saturation state promote rapid lithification and favorable
geochemical conditions for microbialite preservation (Riding, 2005; Riding and Liang, 2005).
The occurrence of the microbialites at the same level as the mass extinction (Mander et al., 2008)
! 131
along with the incredibly widespread distribution of the same growth phases over distances of
10s of km (Ibarra et al., 2014b), indicates the microbialites were able to achieve ecological
dominance over a vast aerial extent due to an impoverished metazoan community.
With respect to carbonate saturation state, the end-Triassic is considered a greenhouse interval
with elevated pCO
2
levels (McElwain et al., 2009; Schaller et al., 2011) induced by CAMP
volcanism. Although high pCO
2
may lead to transient periods of ocean acidification (Greene et
al., 2012b), an increase in pCO
2
has been shown to ultimately result in a long-term period of
carbonate oversaturation (Zachos et al., 2005; Kump et al., 2009; Honisch et al., 2013).
Warming and enhanced silicate weathering increases the flux of calcium carbonate ions to the
oceans creating a preservation mechanism for benthic microbial communities that flourished in
the absence of metazoan pressures. Therefore, while a metazoan post-extinction ‘dead zone’
(Mander et al., 2008) facilitated the development and ecological dominance of extensive
microbial mats, their preservation as delicate microbial mat textures (Fig. 3) and widespread
microfossils (Fig. 4) was largely a product of carbonate oversaturation. A similar scenario has
already been proposed for unusual calcite deposits across the Westbury Formation (Greene et al.,
2012a) and our data indicate subsequent deposition into the Lilstock Formation continued to
favor carbonate deposition up-section.
Our observations are similar to observations made across other times of biotic crisis in
which widespread carbonate microbialite buildups occur in the immediate aftermath of other
extinctions (Mata and Bottjer, 2011). The Cotham Member contains a clear switch from
predominantly skeletal carbonate in the Lower Cotham Member to microbial carbonate in the
Upper Cotham Member across the mass extinction interval (Mander et al., 2008). We would
thus expect other sections across the Triassic-Jurassic interval to contain extensive deposits of
! 132
microbial and/or predominately non-skeletal carbonate in marine and terrestrial settings.
Although the sections of the SW UK have been under scrutiny regarding depositional setting
(Radley et al., 2008), even if the microbialites formed in a normal salinity lagoon (sensu Radley
et al., 2008), their striking lateral extent (Fig. 2A), the extraordinary preservation of intricate mat
textures/fossils (Fig. 3-4), and their δ
13
C
org
values suggest the lagoon was large enough to record
widespread climatic and biogeochemical signals also captured in other Tethyan sections.
Alternatively the occurrence of the microbialites during a transgressive systems tract (Hesselbo
et al., 2004), suggests they were the first organisms to colonize the initial establishment of what
eventually became a shallow carbonate ramp setting.
Furthermore, the Cotham Marble microbialites contain a ‘tubestone’ morphology (Ibarra
et al., 2014b), a unique microbialite form in which microbial mats are spatially arranged into a
mesoscopic honeycomb-like fabric along bedding planes and whose polygonal voids were
subsequently filled with sediment thus resembling ‘tubes’ (sensu Corsetti and Grotzinger, 2005).
Laterally extensive (100’s to >1000’s of km
2
) deposits of tubestone microbialites have only been
described from other ‘greenhouse’ intervals, namely: post ‘snowball Earth’ (Corsetti and
Grotzinger, 2005); Cambrian-Ordovician (Shapiro and Awramik, 2006); Paleocene-Eocene
Thermal Maximum (Lamond and Tapanila, 2003); and similar structures have been described
from the end-Permian (Kershaw et al., 2007). The preservation of widespread ‘tubestone’
microbialites and/or those exhibiting intricate dendritic/thrombolitic fine-grained textures may
result from a pronounced increase in pCO
2
leading to warming and enhanced silicate weathering
(Zachos et al., 2005; Honisch et a., 2013), creating anomalously high carbonate saturation across
shallow water environments ideal for the rapid lithification of delicate microbial mat textures
over vast distances.
! 133
CONCLUSIONS
We present new geochemical and petrographic evidence that indicates the Upper Triassic
‘Cotham Marble’ microbialites formed in response to the events of the end-Triassic mass
extinction. The synchronous occurrence of extraordinarily widespread microbialites at the same
level as (1) the mass extinction, (2) a bloom of prasinophytes, (3) unusually preserved microbial
fossils from nearby intertidal settings and (4) an excursion in δ
13
C
org
that may be related to the
so-called ‘initial carbon isotope excursion’, highlights the non-local environmental mechanisms
that led to their formation. Importantly, the microbialites might represent a microbial carbonate
manifestation of a weathering-induced shallow carbonate burial event that resulted from a
CAMP-related increase in pCO
2
. As seen across other times of environmental crises associated
with episodes of high pCO
2
, the Cotham Marble microbialites reflect a shift in carbonate
deposition from skeletal to microbial across the end-Triassic. The thin nature of the microbialite
unit highlights the need for more high-resolution analyses and microfacies investigations across
this critical time in Earth history.
ACKNOWLEDGEMENTS
We thank Nick Rollins and Will Berelson for assistance with stable isotope analyses of
organic carbon. We also thank Michael Lewis and Ramues Gallois for field assistance and the
Bristol Museum and Art Gallery for access to museum specimens. Field permits for sampling at
Lavenock Point, St. Audries Bay, Lower Woods. This research was supported by the National
Science Foundation (NSF Earth Life Transitions Grant EAR-1338329), the American
Philosophical Society Lewis and Clark Fund for Exploration and Field Research in Astrobiology
to YI, and SEG was additionally supported by UK NERC grant NE/H023852/1s.
! 134
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6725.
Zachos, J.C., Röhl, U., Schellenberg, S.A., Sluijs, A., Hodell, D.A., Kelly, D.C., Thomas, E.,
Nicolo, M., Raffi, I., Lourens, L.J., McCarren, H., Kroon, D., 2005, Rapid acidification
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of the ocean during the Paleocene-Eocene Thermal Maximum: Science, v. 308, p. 1611-
1615.
Zachos, J.C., Dickens, G.R., and Zeebe, R.E., 2008, An early Cenozoic perspective on
greenhouse warming and carbon-cycle dynamics: Nature, v. 451, p. 279-
283,10.1038/nature06588.
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FIGURES
Figure 1. Upper Triassic stratigraphy of the Southwestern United Kingdom. (A) Generalized
stratigraphic column from St. Audries Bay. (B) δ
13
C
org
from St. Audries Bay from (Hesselbo et
al., 2002) and microbialite data from this study (green). (C) Stratigraphic column from Stowey
Quarry and the Mendips region after (Hesselbo et al., 2004). (D) Stratigraphic column from
Lavernock Point.
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Figure 2. Regional Map. (A) Study sites and other Cotham Marble microbialite localities. (B)
Laminated and dendritic clyclic mesostructure of the Cotham Marble microbialites (sample
location: Bristol). (C) δ
13
C
org
of four of the Cotham Marble microbialite layers by location (see
Fig. 2A).
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Figure 3. Cotham Marble microbialite facies. (A) Outcrop photograph of the microbialites from
Lower Woods Natural Reserve (note: the Langport Member does not occur at Lower Woods and
the microbialites are directly overlain by the Blue Lias Formation). (B) Dendritic and
stromatolitic couplet of the microbialites (sample location: Manor Farm). (C) Thin section
photomicrograph of a cluster of round organic-walled microfossils. (D) Close-up of a single
organic-walled microfossil displaying a diagonal suture.
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Figure 4. Microfacies of the Upper Cotham Member from Lavernock Point, South Wales. (A)
Dark carbonate-rich forset lamination under transmitted light. (B) Same field of view as Fig. 4A
under reflected light with arrows denoting filamentous microfossils. (C) Close-up of filamentous
microfossil. (D) Micritic dendrites from the oolitic bed (denoted by white arrows).
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Figure 5. Cross plot of carbonate δ
13
C and δ
18
O of the Cotham Marble.
18
O (VPDB)
13
C (VPDB)
y = -1.2814 + 0.2003x R = 0.29115
Bristol
Manor Farm Quarry
Lower Woods
Stowey Quarry
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Table 1. Stable isotopic compositions of δ
18
O and δ
13
C (‰ VPDB) for carbon and oxygen from
their corresponding site and microbialite layer.
Location Layer δ
13
C
carb
δ
18
O
carb
L1 0.621 -1.241
L1 -0.799 -1.015
L2 -0.479 -1.561
L3 0.576 -0.847
D1 -0.237 -1.548
D2 0.406 -1.026
L1 0.442 -1.119
L1 -0.033 -1.176
D1 -0.146 -2.152
D1 0.078 -0.503
Bristol D1 -0.299 -1.176
L2 0.402 -1.567
L3 0.543 -0.670
D2 -0.472 -1.655
D2 -0.328 -2.028
L1 0.195 -0.458
L2 -0.685 -1.532
L1 0.041 0.021
L1 -1.860 -1.185
D1 0.541 -1.088
L3 -0.191 -1.512
D1 -0.081 -1.852
Manor Farm D1 -1.256 -0.754
L2 -1.710 -2.308
L1 -0.077 -1.926
L1 -0.477 -1.330
Lower Woods L1-d -0.759 -1.701
D1 0.151 -2.730
L1 -2.122 -1.345
Stowey Quarry D1 1.227 -1.207
L2 1.584 -0.633
L2-d 1.213 -0.979
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Table 2. Stable isotopic composition of δ
13
C
org
(‰) of organic carbon of microbialite samples
from their corresponding site and layer.
Location Sample ID Layer % Org C δ
13
C
org
Avg. of
duplicates/
triplicates
MFD1s6 D1 0.16 -28.18
MFD1s6d D1 0.18 -29.03 -28.61
MFL1s2 L1 0.20 -27.03
MFL1s2d L1 0.23 -27.89 -27.46
MFL3 L3 0.18 -27.12
Manor Farm MFL3d L3 0.20 -27.04 -27.08
MFL2 L2 -27.02
MFD2f D2 -27.65
MFD2d D2 -27.34
MFDhed D1 0.233 -28.15
MFDhedd D1 0.237 -27.83 -27.99
MFDL2 L2 0.199 -27.16
MFDL2d L2 0.222 -27.18 -27.17
RGL1 L1 0.21 -29.09
RGL1d L1 0.22 -29.61
Pinhay Bay RGL1d2 L1 0.22 -29.29 -29.33
RGD1 D1 0.23 -28.17
RGD1d D1 0.22 -28.31 -28.24
RGL2 L2 0.16 -27.45
RGL2d L2 0.13 -27.75 -27.60
BL3 L3 0.19 -28.55
BD1s D1 0.22 -27.15
BD1sd D1 0.21 -27.58 -27.37
BD2cL1 L1 -26.71
Bristol BD2cL1d L1 -27.28 -27.00
CMBL2 L2 -26.75
BdL2 L2 0.196 -27.33
BdL2d L2 0.202 -27.23 -27.28
CMBD2f D2 -28.60
CMBD2d D2 -27.29
LWL2 L2 0.16 -27.11
LWdl L1 -26.34
Lower Woods LWwl L1 -25.96
LWf1 D1 0.179 -28.09
LWf1d D1 0.188 -28.02 -28.06
LWf2 D1 0.187 -27.93
LWf2d D1 0.194 -28.07 -28.00
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LWf3 D1 0.199 -27.63
LWf3d D1 0.180 -28.02 -27.82
LWf4 D1 0.205 -27.67
LWf4d D1 0.196 -28.21 -27.94
LWf5 D1 0.186 -28.19
Lower Woods LWf5d D1 0.196 -28.25 -28.22
LWf6 D1 0.176 -28.44
LWf6d D1 0.184 -28.80 -28.62
LWD1hed D1 0.239 -28.60
LWD1hed2 D1 0.230 -28.90 -28.75
LWD2 D2 0.234 -26.86
LWD2d D2 0.217 -27.34 -27.10
LWD2hed D2 0.237 -27.26
LWD2hedd D2 0.222 -27.75 -27.51
L1 -25.8
Stowey Quarry D1 -27.54
D1 -29.36
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Chapter 6:
Lateral continuity of multi-scale stromatolite morphology:
Implications for assessing the dominant scales of control
ABSTRACT
A key focus in the study of ancient stromatolites and other potential microbialites is to
determine which processes control their morphology, essential for biosignature detection and
paleoenvironmental reconstruction. The process(es) that govern the formation of stromatolites
occur at various scales (local versus regional), however, determining the relative importance of
local versus regional processes of control remains a challenge, particularly for ancient deposits.
We build upon the traditional multi-scale level approach of morphological investigations by
including a lateral continuity component of fine- to large-scale features to determine the relative
importance of local versus non-local controls on stromatolite morphology. Observations on
stromatolites from the Miocene Barstow Formation, and marine microbialites from the upper
Triassic Southwestern United Kingdom reveal the following: (1) lateral patterns are strongly
dependent on the nature of the depositional system, (2) some features are controlled by local
processes whereas others are primarily controlled by regional processes, (3) in some instances,
the scale of the feature does not necessarily scale with the scope of the process. The combination
of lateral continuity in the context of hierarchical morphological properties can provide an
effective approach for gauging the dominant controls and spatiotemporal significance of
stromatolite texture and morphology.
INTRODUCTION
Lamination is considered one of the principal features of stromatolites (Kalkowsky,
1908)—lithified, sedimentary growth structures commonly regarded as Earth’s earliest
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macroscopic fossils (Hofmann et al., 1999; Allwood et al., 2006). However, the detailed
mechanisms that govern the formation, lithification, and diagenesis of stromatolite laminae
across the rock record remain obscure (Grotzinger and Knoll, 1999). Biological, chemical, and
physical processes are considered to play key roles on stromatolite formation (Cloud and
Semikhatov, 1969; Awramik and Riding, 1988), yet these processes could be embedded in the
fabric at various scales in the form of micro-meter to m-scale features and textures, making it a
challenge to resolve their relative role and importance on stromatolite laminae and overall
stromatolite morphology. Studies of modern microbial mats have highlighted several processes
that may lead to the formation of various microbialite textures across observable timescales
(Dupraz et al., 2006). However, direct comparisons to ancient stromatolites with respect to
lithification, texture, and the role of microorganisms have raised concern about the degree to
which modern microbial mat investigations relate to ancient lithified stromatolites (see review in
Grotzinger and Knoll, 1999). Given the potential ecological significance of stromatolites across
the geological record, it is important to assess the morphological and geochemical imprint of the
various local and nonlocal processes responsible for their formation (e.g., Bosak et al., 2009;
Kah et al., 2009; Mata et al., 2012).
With improved technology and geochemical capabilities, there has been increasing
interest on sub-m-scale features with a particular rise in microscopic, high-resolution
investigations (e.g., Reitner et al., 2011). Such focus on sub m-scale features (e.g., laminae, m-
scale shapes, porosity, branching, microfossils, etc.) has preferentially drawn attention to sub–m-
scale local controls (e.g., the role of microbes, diagenetic alteration, local hydrodynamics, high-
resolution geochemical changes, etc.), potentially masking the role of large-scale, non-local
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controls on the textures in question, and possibly resulting in misleading or partial
interpretations.
To improve interpretations on observations of ancient stromatolites, this study builds
upon the traditional multi-scale approach (Shapiro, 2000) by incorporating a lateral continuity
component to provide a method of assessment of the relative importance of scale and dominant
controls that may contribute to stromatolite morphogenesis. A multi-scale, high-resolution,
lateral continuity analyses of stromatolite textures from the Miocene Barstow Formation
lacustrine stromatolites and marginal marine stromatolites from the Upper Triassic beds of the
Southwest United Kingdom is presented. We focus on banded deposits as they are (1) most
representative of deposits found throughout most of Earth history (Riding, 2011) and (2) deposits
that are increasingly difficult to reproduce in the laboratory and rarely found lithifying in modern
environments, warranting alternate complementary methods of investigation. The analysis of
multiple specimens from laterally continuous sections may allow us to disentangle the dominant
scales of control on stromatolite laminae and overall morphology, and in turn potentially
improve our interpretations of their environmental significance across the geological record.
PROCESSESS AND SCALES OF CONTROL
In most depositional systems, morphological attributes of stromatolites are controlled by
changes in organic, inorganic, and post-depositional processes, each of which may act at various
scales and exert varying degrees of control on the fabric (Table 1). Micro-scale processes that
can impart local textural changes include crystal growth/dissolution, neomorphism (diagenesis),
and microbial processes via the presence/absence of the microbial cells and colonies by inducing
crystal growth and/or passively altering the micro-texture (Dupraz et al., 2009). These
microscopic observations, which can be made via petrographic analyses (light microscopy,
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scanning electron microscopy, etc.), are usually hierarchical in nature. That is, most micro-meter
scale processes can be scaled to the mm- and cm-scale; microorganisms can form colonies and
biofilms to increase their influence on the accreting substrate, and similarly, although crystals
nucleate at the micro- to sub micro-meter scale, they may grow to the cm-scale or larger.
Therefore, observations made at the micro-scale can benefit from petrographic analyses that
expand the field of view to the mm and cm-scale (e.g., photomosaic petrography) to reveal mm
and cm-scale textural relationships with highly resolved microscopic elements. Mm-scale
processes constitute grain incorporation into the fabric (e.g., trapping of grains), crystal growth,
consortia of bacteria and biofilms. Similarly, diagenetic processes like dissolution,
neomorphism, and weathering may alter the fabric at various scales and additionally contain a
critical temporal component, potentially acting upon stromatolite micro- to macro-scale features
at any time after deposition.
Relatively larger-scale processes that may contribute to stromatolite fabrics include those
that exert their control at the m-scale scale or greater (Table 1). Like local processes, non-local
processes vary strongly by setting. In lakes and marine environments, for example, currents and
wave activity, usually dominant at the scale of m’s to 10’s of meters, may alter stromatolite
laminae and cause the structures to be deformed before they are fully lithified. These m-scale
processes are also likely responsible for the m-scale shapes of biostromes, possibly controlling
their m-scale spatial distribution (e.g., Hoffman, 1976; Andres and Reid, 2006; Bosak et al.,
2013b). In fluvial spring settings, however, the morphology of stromatolites and other associated
precipitates is dominantly controlled by lateral gradients of flow that cause sub m-scale lateral
variations in temperature, nutrients, turbulence, etc., that are continuously altered downstream
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from the source vent resulting in highly distinct and diverse morphologies amongst adjacent
deposits (Fouke, 2000).
Finally, there are larger-scale processes 100’s of m to km-scale that can be captured by
stromatolites and ultimately influence sub-meter scale textures, like changes in climate and
seasonal changes in water chemistry that may cause biogeochemical changes over vast distances.
It is well-recognized, for example, that in environments where the local controls may be subdued
(e.g., closed basin lakes), given enough external control, stromatolites can record high-frequency
climatic changes capturing region-wide, if not global signals (Benson et al, 1990; Berelson et al.,
2009; Petryshyn et al., 2012). In order to establish a clear idea of the scope of the significance of
specific sub m-scale features, however, it is important to get a relative sense, if not determine,
which processes (local or nonlocal) employed the greatest degree of control on the textures being
investigated.
LATERAL CONTINUITY
Although the formation of stromatolite laminae is the result of a complex mix of various
processes (Table 1), it may be possible to dissect the dominant scales of control on stromatolite
textures by measuring the degree to which various textures vary laterally along strike.
Ecological interactions have been shown to decrease spatially with increasing geographical
distance (Nekola and White, 1999), a process known as distance decay. Similar observations can
be expected for two stromatolite deposits forming at the same time in a shared depositional
system (Fig. 1), such that the greater the distance between two deposits, the higher the likelihood
of physical, chemical, or biological chaos (noise or randomness) between the two, thus resulting
in a higher likelihood for morphological differences (e.g., Pope and Grotzinger, 2000). We
know however, that there are stromatolite-forming systems in which the opposite is observed and
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adjacent deposits can contain vastly different morphologies (Logan et al., 1964) and likewise,
distant deposits can result in nearly identical textures (Sumner, 1997). In cases where lateral
continuity is reduced between adjacent deposits, we can assume that local controls are dominant
(Table 1). Conversely, in instances where lateral continuity of textures can be easily traced, it is
likely that those traceable features are the product of processes whose scale of control extended
for at least the minimum distance between the two deposits. Indeed, previous investigators have
utilized the concept of ‘lateral linkage’ between stromatolites and lateral continuity of structures
across large and small distances (Logan et al., 1964; Sumner, 1997; Pope and Grotzinger, 2000;
Allwood et al., 2007; Kah et al., 2009; Van Kronendok, 2011) to infer local or environmental
controls on morphology. In particular, lateral continuity of stromatolitic domes has been applied
to studies of basin analyses (Surdam and Wray, 1976). However, the use of lateral continuity to
determine possible dominant controls has not been explicitly presented, nor has this method been
combined with multi-scale investigations. Although this method does not reveal explicit
processes of control it can be used to constrain the scale of the significant processes of control
and in turn inform our understanding of the environmental scope of the textures under
investigation.
Below we present two examples in which we examine the degree to which stromatolite
textures vary laterally along strike. We build upon the traditional multi-scale approach (macro-,
meso-, micro- sensu Shapiro, 2000) and expand the scale of the investigation to include analyses
of the lateral continuity of fine- to large-scale morphological attributes. It is important to note
that all of these examples focus on systems dominated by precipitation as opposed to
sedimentation. Systems dominated by sedimentation will inherently result in a high degree of
variability among adjacent mounds (Fig. 1).
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CASE STUDIES
Barstow Formation Lacustrine Tufa
The Barstow Formation outcrops about 13 km north of Barstow in the Central Mojave
Desert San Bernardino County, southern California. It is about 1000 m thick and consists of a
fining upward succession of basin-fill deposits of alluvial, fluvial, and lacustrine origin (Dibblee,
1968; Woodburne et al., 1990). The middle and upper members consist of localized calcium
carbonate mounds referred to as ‘tufa’ among predominantly sandstone and claystone layers.
Here we focus on a carbonate unit (~8 m thick) that occurs at the base of the Middle Member.
The carbonates occur as m-scale mounds and are laterally discontinuous.
Tufa deposits from the south limb of the Owl Campground locality occur as 20 to 1.5 m
thick mounds and towers or 2 to 1.5 m thick beds. The mounds overlie carbonate cemented
vertical pipes that display vertical and concentric growth. Stromatolitic tufa samples described
in this study include decimeter-scale subunits of large, localized meter-scale mounds (Fig. 2).
The subunits are cylindrical/nodular in shape and precipitated vertically relative to the growth
surface (Fig. 2). Isopachous banding developed around the object of nucleation—possibly a root
or stem from aquatic plants (Pedone and Caceres, 2002). Vertical cross sections of samples
illustrate two dominant mesoscopic textures: dense banded fibrous calcite and porous weakly
laminated calcite (Fig. 3). Microscopically, the banded calcite is composed of fibrous spar layers
(~100 – 500 µm thick) separated by micrite layers (~50 µm thick) of variable thicknesses.
Fibrous spar layers are laterally persistent across subunits and maintain a constant thickness
whereas the micrite layers vary in thickness along individual laminae (Fig. 4)
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Interpretation of scales of control on stromatolite fabric
At the macroscale tufa mounds of the Barstow Formation formed when Ca-rich
groundwater entered he alkaline saline lake via subaqueous springs. The bicarbonate-rich lake
water would have fostered localized carbonate formation at sites where groundwater entered the
lake (Park 1995; Cole et al., 2004; Becker et al., 2001). Localization of stromatolitic tufa
deposits around springs is a common formation mechanism observed in many modern and
ancient lacustrine deposits (Scholl, 1960; Bischoff et al., 1993; Arp et al., 1999). At the
mesoscale, the occurrence of tubular voids and tubular secondarily filled molds suggests aquatic
plants served as suitable sites for nucleation as recognized by Pedone and Caceres (2002).
Nucleation around aquatic plants is common in other lacustrine carbonates (Arenas et al., 2000;
Riding, 1987). Mesoscopic laterally continuous dense bands of constant thickness reflect
carbonate precipitation controlled at the scale of the m-scale mound or greater (spring activity,
lake chemistry, climate). On the other hand, uneven micrite laminations are more strongly
controlled by local cm-mm mechanisms (microbial activity, cement growth) (Fig. 4).
Upper Triassic Cotham Marble
The Upper Triassic Cotham Marble (CM) microbialites are a ~20 cm thick deposit that
occur as m-scale mounds and extends for over 2,000 km
2
across the southwestern United
Kingdom (SW UK) (Fig. 5). The deposits formed during the Triassic-Jurassic transition, a time
of significant biotic and environmental change (Raup and Sepkoski, 1982). Detailed mesoscopic
investigations of the CM from several locations reveal a cyclic pattern of sedimentation
composed of alternating laminated and dendrolitic mesofabrics whose distinct layers can be
traced across distances of over 20 km (Ibarra et al., 2014b). Microscopically, laminated layers
are dominantly micritic. Dendrolitic layers are composed of distinct branching microbial mat
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textures with well-preserved evenly spaced branching patterns (Ibarra et al., 2014b). The
dendrolites are micritic and contain a distinct dark outline. Pyrite-coated filamentous putative
microfossils are found restricted to the dendrolitic fabric suggesting a biogenic origin for the
dendrolites. Remarkably, individual sub-mm laminae in some of the laminated layers can be
traced from site to site over distances of at least ~100 km (Fig. 5).
Interpretation of scales of control
Macroscopically, the Cotham Marble microbialites are shaped as discrete dm to m-scale
domes. Meter-scale processes like currents and water movement likely control variability of
intermound shapes and sizes. Indeed, a flat-pebble conglomerate is often reported to occur
between adjacent mounds (Hamilton, 1961) suggesting the presence of channels or currents that
may have shaped the distances between adjacent mounds (e.g., Hoffman, 1976).
Mesoscopically, the changes from laminated to dendritic can be traced over vast distances thus
reflecting widespread cyclic non-local controls (changes in climate, water chemistry, or
seasonality across the various locations). At the sub-mm level, laminated layers reflect controls
at the mega-scale given how well they can be traced from mound to mound and across the
various locations (Fig. 5). Instances where individual laminae are truncated or vary in thickness
along a single band indicate localized controls on the texture.
The dendrolite layers reflect a unique interplay between local and nonlocal controls
making it difficult to determine which processes (local or nonlocal) impose the greatest
dominance on their morphology. On the one hand, dendritic layers can be traced laterally for
tens of kms reflecting regional controls on their morphology. On the other hand, each
dendrolitic structure is morphologically unique despite displaying similar dendrolitic and evenly
spaced branching patterns throughout the sites investigated thus highlighting the strong role of
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local controls on micro-scale variability (Fig. 5). In this unique case, given the excellent
preservation of microbial mat textures and putative microfossils, it is likely the local controls
strongly reflect unique branching patters of microbial mats. Therefore, embedded within the
broader scope of a strong environmental control, microbial mats are able to exert control of the
resulting sub-cm-scale fabrics suggesting that the dendrolitic structure itself is a product of
strong microbial control (as microfossils suggest), but the widespread presence of the microbial
mats ultimately result from nonlocal controls.
DISCUSSION OF RESULTS FROM THIS STUDY
Lacustrine and marine stromatolite examples from this study contain banded textures that
were common in marine environments for over 2 billion years of Earth’s history and continue to
be prevalent in modern lacustrine and spring settings today (Riding, 2011). Analyses of lateral
continuity of sparry, micritic, and dendrolitic mesofabrics in the examples above demonstrate the
utility of the lateral continuity approach in analyses of the relative controls on ancient
stromatolites. While the use of observations and analyses that focus on sub-m scale features are
important for assessing the relative control of local processes (Table 1), considering the potential
role of non-local processes may ultimately reveal the degree to which the textures under
investigation reflect local vs. non local processes. Specifically, this study presents the following
observations:
(1) Relatively high-frequency changes of dominant controls can occur at small (cm-scale)
scales (Fig. 4). For instance, lateral variability of textures for the Barstow tufa change at least
three times over the scale of about 5 cm (Fig. 4) reflecting relatively high resolution textural
changes in the dominant processes of control and thus highlighting the importance of high-
resolution analyses in attempts to infer environmental relevance.
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(2) The size of the features observed do not necessarily scale with the extent of the
dominant processes that control those features. For example, one of the most striking
observations of the Barstow tufa subunits is the degree to which mesoscopic textures can be
traced laterally (Fig. 4) despite the unique macro-morphology of each subunit (Fig. 3). However,
at the mesoscale, the process responsible for their overall macroscopic shape is nucleation
around aquatic plants, a highly localized phenomenon leaving a conspicuous imprint on the
successive fabric. Conversely, the sparry laminae of the Barstow tufa—and similarly the
micritic laminae of the CM—are relatively small, on the order of 1 mm or less, yet the dominant
processes responsible for their formation are more wide-reaching. This observation highlights
that the more wide-reaching significant growth variables may not always translate into highly
conspicuous textures and may consequently go unnoticed when investigating a single subunit or
at coarse resolution.
(3) This study presents an example of how lateral continuity can help address how
processes at various scales interact to produce the observed features/textures in stromatolites.
The dendrolitic layers of the CM resulted from the interaction of highly localized and non-
localized controls. Without the lateral continuity analyses and given the presence of putative
microfossils within the dendrolitic fabric (Ibarra et al., 2014), local evidence would suggest
microbial controls are the dominant process in control of the dendrolitic fabric. However, the
expansive continuity of the same dendrolitic fabrics over a span of over 100 km indicates the
non-local processes (e.g., water chemistry, depth, climate) were the main drivers of their
widespread occurrence. The dendrolitic fabric—interpreted as microbial in origin—is in fact
embedded within a larger, wide-reaching variable of control that ultimately determines its
development.
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(4) Potential setbacks associated with this method include the fact that laterally
continuous sections may not always be available due to tectonics or other post-depositional
processes. Also, in some instances, it may be difficult to find time-correlative samples if the
beds cannot easily be traced along strike. Furthermore, this method is most sensitive to
stromatolites that formed via in situ precipitation, as stromatolites influenced strongly by
sedimentation of grains (e.g., trapping and binding) will inherently have relatively higher
variability along adjacent deposits resulting in textures that preferentially reflect local processes
(Awramik and Riding, 1988).
BROADER IMPLICATIONS/SIGNIFICANCE
A major unresolved issue in the study of ancient stromatolites is determining the
dominant processes responsible for their morphology, critical to precisely discern
paleoenvironmental information. Given that stromatolites comprise a diverse range of
morphologies from macroscale mounds, towers, domes, to mesoscopic banding, lamination,
branching, to microscopic textures and mineralogy—different processes over various temporal
and spatial scales contribute to the resultant morphology (Table 1). However, in stromatolite
studies, to date, the combined use of multi-scale descriptions together with lateral continuity
investigations is not readily applied.
Given that different processes impose their effects on morphology at different scales the
combination of both methods may prove an essential and promising step towards unraveling the
challenging task of determining dominant controls. Scaling issues are widely recognized in
studies of processes and the development of geomorphic landscapes, and it is understood that
they are governed by a range of processes that operate at a range of scales (de Boer, 1992).
Unlike other systems in environmental science that utilize spatial relationships to study
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hierarchical patterns of control (e.g., Caron et al., 2008), stromatolites, by definition as lithified
structures may retain morphological attributes of those controls, many of which could be
preserved for billions of years (Allwood et al., 2009). Therefore, the use of lateral continuity
may allow for characterization of the nature of spatial variation to more precisely interpolate the
significance of measurements made at various scales.
As measurements on stromatolite textures continue towards increasingly localized and
microscopic techniques, it is important to be able to upscale fine-scale measurements and
extrapolate the full scope of their potential environmental significance. Relating fine-scale
stromatolite patterns (which we are arguably the most familiar with) to broad scale phenomena
and controls wherever applicable remains a critical challenge, yet is essential to being able to
understand their potential as paleoecological archives and astrobiological targets (Cady et al.,
2003). Furthermore, a general understanding of regional spatial controls using high-resolution
lateral continuity may help reduce the number of factors that must be investigated and also
inform at which scales measurements should be done (e.g., Andres and Reid, 2006; Planavsky
and Ginsburg, 2009).
Finally this study shows that when investigated under the framework of combined lateral
continuity and multi-scale observations, multiple processes (local vs. non local) may be
responsible for controls at various scales and in most instances determining which specific
controls are dominant is not straightforward (Awramik and Grey, 2005). Most ancient
stromatolites have experienced varying degrees of diagenesis making it difficult to decipher the
relative contribution of primary versus altered fabrics (Grotzinger and Knoll, 1999). Therefore,
while lateral continuity may be used to infer a dominance of relatively local versus nonlocal
controls, it is not a means of assessing specific growth variables. If a specific process is invoked,
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other more conclusive methods should be used to assess controls (Allwood et al., 2009). For
instance, high variability between adjacent stromatolites or even wrinkly laminae at the sub-mm
scale does not necessarily invoke biogenicity but rather, more broadly reflects highly localized
controls on variability (Table 1).
CONCLUSIONS
A combined multi-scale and lateral continuity analysis is presented for banded lacustrine
stromatolites and marginal marine stromatolites to assess the potential relationships between
growth variables and multi-scale stromatolite fabrics. Relatively high-resolution analyses of
laterally continuous features reveal that (1) lateral patterns are strongly dependent on the nature
of the depositional system, (2) some features are controlled by local processes whereas others are
primarily controlled by regional processes, and (3) in some instances, the scale of the feature
does not necessarily scale with the scope of the process. These results have implications
regarding one of the most longstanding enigmas in the study of ancient stromatolites –
determining the relative role of various processes on stromatolite laminae and their overall
morphology. The combination of these approaches thus far remains an underutilized tool for
assessing the relative importance of stromatolite textures and for assisting in determining the
significance of stromatolites through space and time.
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FIGURES
Figure 1. Expected lateral continuity relationship between two stromatolite subunits according
to distance decay. Systems generally follow the trend in the line except for when local or non-
local processes dominate. After Nekola and White, 1999.
Relative distance between stromatolite units
Lateral Continuity
continuous discontinuous
near far
LOCAL
CONTROLS
NON-LOCAL
CONTROLS
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Figure 2. In situ macro and mesostructure of Barstow Formation tufa. (A) Carbonate mound.
(B) subunit showing concentric growth pattern.
B A
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Figure 3. Mesostructure of several Barstow Formation subunits. (A-C) Mesoscopic morphology
of tufa subunits, note the distinct shapes. (D-F) Cross-sections of units A-C displaying strikingly
similar internal growth phases.
1 cm 1 cm 1 cm
1 cm
1 cm
1 cm
A B C
D E
F
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Figure 4. Thin section photomicrographs of Barstow Formation subunits. (A-C) Microstructure
of three different subunits (see Figure 3) highlighting three dominant textural phases: porous-
weakly laminated, micrite laminea, and fibrous spar laminae.
Spar
laminae
Micrite
laminae
Porous,
weakly
laminated
Texture Variability across subunits Scales of control
spar laminae low m-scale control (e.g., spring
activity, lake chemistry)
micrite laminae moderate m-scale through micro-scale
porous laminae high micro-scale controls (microbes,
cement growth, dissolution)
C B A
1 mm 1 mm 1 mm
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Figure 5. Upper Triassic Cotham Marble microbialites. (A) Map displaying known Cotham
Marble microbialite localities. After Mander et al., 2008. (B-D) Mesostructure of dendrolitic
and laminated phases of the Cotham Marble from Bristol, Lower Woods, and Pinhay Bay,
respectively.
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Table 1. Potential processes of control on stromatolite morphology at various scales.
Scale of
control
Morphogenesis
(inorganic)
Biogenesis
(organic)
Diagenesis
(post-deposition)
Local
µm
crystal growth;
local water chemistry
microbes cementation,
dissolution,
precipitation;
replacement/inversion;
micritization
mm
crystal growth;
sedimentation
microbial colonies;
biopolymers;
passively altering
textures; actively
inducing
precipitation
cement, dissolution,
precipitation
cm
crystal growth; cement
growth; sedimentation
biofilms; microbial
mat development;
trapping and binding
of sediment
cementation,
dissolution,
precipitation
m
spring outflow, lateral
gradients of (physical,
chemical) flow,
turbulence (currents,
waves, tides) turbidity
biofilms, microbial
mats; microbial
chemical/nutrient
recycling
soft sediment
deformation; burial;
slumping; desiccation;
displacive growth by
evaporites
Non-
local
km
climate, seasonality,
water level/depth,
water chemistry,
nutrient supply
widespread
microbial dispersal
controlled by other
factors (see km-scale
morphogenesis)
burial; compaction;
tectonic deformation
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Chapter 7:
Multiscale controls on the formation of lacustrine tufa from the
Middle Miocene Barstow Formation, California
ABSTRACT
Meter-scale tufa (carbonate) mounds of the lacustrine sediments of the Miocene Barstow
Formation reveal multi-scale and multi-process controls on tufa formation. Tufa samples
examined here include several decimeter-scale nodular sub-units that originate from larger
meter-scale mounds. The subunits display tubular (~2 mm in diameter) molds interpreted as
growth around aquatic plants. Like many lacustrine carbonates, the samples are strongly banded
on the cm/mm scale; mesofabrics alternate between dense banded carbonate and porous weakly
laminated carbonate.
Microscopically, the dense banded carbonate is characterized by (1) laterally continuous
fibrous spar/micrite couplets and (2) micritic laminae characterized by an irregularly laminated
clotted texture. The porous regions consist of microclots, peloids, and filamentous microfossils
in a matrix of microspar. The spar laminae can be traced unambiguously from subunit to sub-
unit, whereas the micritic fabrics do not maintain a constant thickness across sub-units.
The tufa mounds of the Barstow Formation are products of multi-scale controls that each
left behind a unique imprint on their morphology: m-scale localization caused by spring activity,
cm-scale sites for nucleation provided by aquatic plants, and sub-mm-scale changes in texture.
The occurrence of filamentous microfossils, diatoms, fenestral fabric and a ubiquitous clotted
micritic texture present evidence for a significant microbial component. Interestingly, the
micritic textures vary in thickness from sub-unit to sub-unit, suggesting microbial influence was
locally controlled at the cm-scale. In contrast, the spar laminae maintain constant thickness
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between sub-units across the entire tufa mound, suggesting control at or beyond the m-scale
mound, likely representing larger scale changes in lake chemistry, climate, and/or spring activity.
Given the common use of lacustrine carbonates as environmental archives and partial
analogs of early Earth microbialites, conducting multi-scale level of analyses allows us to more
thoroughly evaluate the different factors that control their formation. Importantly, the scale of
the feature does not necessarily scale with the scope of the process.
INTRODUCTION
Lacustrine tufa mounds are calcium carbonate deposits that result from organic and
inorganic processes at sites where calcium-rich groundwater mixes with alkaline lake standing
water (Scholl and Taft, 1964; Kempe et al., 1991; Arp et al., 1999b; Rosen et al., 2004).
Geochemical analyses on modern and ancient lacustrine tufa deposits have been used to
reconstruct lake histories, serving as records of local climate (Li et al., 2008; Nehza et al., 2009),
lake level fluctuations (Cohen et al., 1997; Hart et al., 2004; Berelson et al., 2009) and as
geochronological tools (Newton and Grossman, 1988; Szabo et al., 1996; Cole et al., 2005).
Lacustrine tufa deposits have also received a lot of attention for their morphological
resemblance to Precambrian microbialites ( Laval et al., 2000; Kazmierczak and Kempe, 2006),
and have furthermore been used to test the theory of an alkaline ocean early in Earth history
(Kempe and Degens, 1985; Arp et al., 1998; Kazmierczak et al., 2011). As a result, modern
lakes of high alkalinity have been a key target for investigations of microscale processes (e.g.,
microbial communities, metabolisms, and chemistry) associated with carbonate precipitation
around microbial mats and biofilms (Thompson et al., 1990; Kempe et al., 1991; Ferris et al.,
1997; Rosen et al., 2004; Gischler et al., 2008; Last et al., 2010).
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However, diagenetic processes often complicate our ability to differentiate between the
possible ranges of environmental factors that control the formation of ancient carbonate deposits.
Despite these limitations, given increasing use of lacustrine carbonate deposits as environmental
indicators and as partial analogs of some of the earliest records of life on Earth, it is critical to
thoroughly investigate the possible range of processes that control their formation and assess
how these mechanisms might translate into their morphology and texture. It is also important to
evaluate the biologic integrity of microfabrics that correspond to frequently used mesoscopic
signatures of life (e.g., stromatolitic, thrombolitic, dendritic) especially in light of studies that
demonstrate dominantly abiotic mechanisms (or lack of unequivocal biologic influence) in the
formation of similar deposits (Jones and Renaut, 1995; Pope and Grotzinger, 2000; Pentecost,
2005; Petryshyn and Corsetti, 2011). Conducting multiscale analyses and investigating potential
diagenetic overprints of tufa deposits serves to differentiate between the different scales of
control on their formation, allowing us to better interpret their implications in modern and
ancient settings.
The lacustrine sediments of the Middle Miocene Barstow Formation contain m-scale tufa
mounds composed of nodular and domal calcium carbonate that formed in a saline alkaline lake.
A unique uranium enrichment of the tufa deposits allowed for U-Pb dating of the deposits (Cole
et al., 2005). Their morphologic and petrographic resemblance to tufa mounds that contain clear
biologic influence, led previous workers of the tufas to infer biologic/microbial influence (Cole
et al., 2004; Cole et al., 2005). However, an in-depth investigation of microscale influences on
carbonate formation has not yet been conducted.
The focus of this work is to critically evaluate the Barstow Formation tufa deposits for
microscale controls on their morphology and texture. We combine meso- and microfabric
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analyses with previously reported megastructure and geomorphological observations to assess
the range of possible controls on tufa formation.
GEOLOGIC SETTING
The Miocene Barstow Formation (~1,000 m thick) outcrops approximately 13 km north
of Barstow in the Mud Hills of the Central Mojave Desert, San Bernardino County Southern
California (Fig. 1). The formation consists of a fining upward succession of basin-fill deposits of
alluvial, fluvial, and lacustrine origin (Dibblee, 1968; Woodburne et al., 1990). Following
deposition, strike-slip motion and related folding formed the Barstow Syncline (Ingersoll et al.,
1996).
The Barstow Formation is divided into three members, the lower Owl Conglomerate
Member, and the middle and upper members (Woodburne et al., 1990). The Owl Conglomerate
Member is composed of coarse fanglomerate of granitic rocks derived from highland sources
north and south of the basin (Dibblee, 1968). The middle and upper members consist of
sandstones, claystones with interbedded ash layers, and localized limestone beds and nodules
(referred to as tufa). Sandstone facies are interpreted as alluvial outwash on the lower floodplain
and the claystone-shale facies and carbonate layers are interpreted to have formed in a large
shallow lake (Dibblee, 1968). Sedimentary and geochemical evidence from arthropod-bearing
concretions located in the middle member indicate that paleolake Barstow was a shallow, saline-
alkaline, and poorly mixed system (Park, 1995; Park and Downing, 2001).
Here we focus on the previously noted “algal limestone” unit (~ 8 m thick) that occurs at
the base of the middle member, directly overlying the Owl Conglomerate Member (Dibblee,
1968). The tufas occur as m-scale mounds and are laterally discontinuous. It has been suggested
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that calcite precipitated into mounds when Ca-rich spring water seeped into the lake (Cole et al.,
2005; Park, 1995).
PREVIOUS WORK ON THE BARSTOW TUFAS
Megascale Distribution of Tufas
Calcareous tufa layers have been reported from three main localities: the phytoherm
locality, a tufa locality on the north limb of the syncline, and the Owl Campground tufa site on
the south limb of the syncline (Fig. 1; described in Cole et al. 2004). At the Owl Campground
site (the focus of this study), the tufas occur as 20 cm to 1.5 m thick mounds and towers or 2 cm
to 1.5 m thick beds. The mounds overlie carbonate cemented vertical pipes that display hollow
centers and concentric growth (Becker et al., 2001). The localized development of the mounds
and presence of the pipes suggest that the pipes could be areas where groundwater might have
entered the lake (Becker et al., 2001).
At the phytoherm locality, a 1–3 m thick tufa unit formed upon the Owl Conglomerate
Member and marks the base of the middle member. Large (3–5 m in diameter) elliptical to
circular tufa mounds occur discontinuously throughout the horizon. The observation of 0.3 to 1
mm diameter stem or branch molds of putative aquatic plants led Pedone and Caceres (2002) to
interpret carbonate mounds at this locality as ‘biolithite mounds’ formed by the calcification of
macrophytes. A third tufa locality has been described on the north limb of the syncline but its
stratigraphic position is obscured by Quaternary alluvium (Cole et al., 2004; Cole et al., 2005).
Geochemical Analyses and Dating
Petrographic analyses of the tufas revealed two dominant fabrics: (1) densely laminated
calcite cement and (2) a spongy nonlaminated micritic fabric (Cole et al. 2004). The fibrous
calcite exhibits elevated levels of uranium whose cyclic variations might reflect changes in the
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U/Ca ratio due to episodic mixing between Ca-rich sprig water with a low U/Ca ratio and Ca-
poor saline alkaline lake water with a high U/Ca ratio (Becker et al., 2001). U-Pb dating of five
tufa samples resulted in ages (2σ) of 14.81 ± 0.39 Ma, 15.30 ± 0.25 Ma, 15.39 ± 0.15 Ma, 16.14
± 0.40 Ma, and 16.24 ± 0.23 Ma (Cole et al. 2005). These ages are in agreement with previously
established Ar-based geochronology of interbedded ash layers.
METHODS
Sample preparation
Calcareous tufa samples examined in this study were collected from the south limb of the
Barstow syncline at the Owl Canyon Campground tufa-bearing site (Fig. 1). Observations on
tufa samples were conducted at the macro-, meso-, and microstructural scale as illustrated by
Shapiro (2000). Tufa samples were slabbed, polished, and scanned on flat bed scanner in order
to observe internal structures and laminae at a mesostructural scale. Representative thin sections
of each sample were examined via light microscopy for microstructural features of the carbonate
fabric. Stable isotopic measurements of carbonate carbon and oxygen were performed in the
Stott lab at the University of Southern California. All of the investigated samples were micro-
drilled at high resolution to target specific textures (e.g, banded spar, banded micrite).
Stable isotopic analyses
Isotopic analyses were conducted on an Elementar Americas Inc. (Micromass Ltd)
Isoprime stable isotope ratio mass spectrometer (IRMS). Samples were measured relative to
CO
2
reference gas calibrated against the NBS-19 (δ
18
O value +2.20‰, δ
13
C value +1.95‰)
carbonate standard, which allows for normalization to the 2-point VPDB-LVSEC isotopic scale.
The precision of this determination is better than 0.06‰ and 0.04‰ (1σ, n = 20) for δ
18
O and
δ
13
C, respectively. A working standard carbonate, (δ
18
O −1.88‰, δ
13
C value of + 2.07‰)
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monitors precision during the course of the run to 0.06‰ and 0.06‰ (1σ, n = 5) for δ
18
O and
δ
13
C, respectively.
RESULTS
Mesostructure
Tufa samples described in this study include several decimeter-scale subunits of large,
localized meter-scale mounds (Fig. 2). The subunits are cylindrical/nodular in shape and
precipitated vertically relative to the growth surface (Fig. 3). Some of the tufa samples exhibit
thin (~2 mm diameter), hollow and secondarily infilled tubular internal molds (Figs. 3C–D).
Isopachous banding developed around the once existing object of nucleation—possibly a root or
stem from aquatic plants (as suggested by Pedone and Caceres, 2002). However, such growth is
not a requirement for the development of the tufas, as some samples do not appear to contain
central nucleation sites (Fig. 3E). A vertically-cut, polished slab and corresponding
photomicrograph illustrates the two most common internal textures: dense banded carbonate and
porous weakly laminated carbonate (Fig. 4). All samples investigated in this study display
alternating episodes of the two dominant mesofabrics.
Microstructure
Microscopic investigations were performed with an emphasis on the porous weakly
laminated carbonate and the banded carbonate detected in the mesostructure of the tufa subunits
(Fig. 4).
Microstructure of Porous Weakly Laminated Carbonate
In thin section, the porous weakly laminated carbonate consists of a porous framework of
fine-grain calcite, largely microspar and micrite (Figs. 5A–C). Some areas contain micritic
peloidal aggregates that interlock in a matrix of spar (Fig. 5B). Spar crystals typically radiate
! 182
from clusters of peloidal, opaque microclots (~50–100 µm in diameter) (Figs. 5B–C). In other
areas the porous regions consist almost entirely of indistinctly clotted micrite (Fig. 5C). Micritic
shrubs and micritic fabrics with distinct filamentous molds were also detected (Figs. 5D–E).
The porous regions are weakly laminated due to discontinuous irregular laminations
(~100 µm thick) composed of fibrous spar and micrite couplets (Figs. 5F–G). Isopachous but
discontinuous fibrous spar laminae developed on micrite layers that vary in thickness within
individual bands. The fibrous spar laminations exhibit a radial fibrous texture that nucleated on
thicker portions of the micrite laminations (Fig. 5D). Laminations of clotted peloidal micrite
also contain laminoid fenestra of irregular shapes (Fig. 6).
Oriented filamentous microfossils were found in highest concentration in the porous
regions (Fig. 7). The calcified filaments are approximately 10 µm in diameter and range from 50
µm to almost 200 µm in length. Filament fragments are oriented parallel to one another and
form distinct layers (Fig. 7). Some filaments are preserved as molds in a matrix of spar and
others were calcified and preserved in a matrix of cloudy micrite. In some cases, a dense micrite
fabric with embedded filaments serves as the primary framework of the microstructure (Fig. 8).
A cross section through a secondarily filled void highlights a clotted micrite fabric that envelops
the sub-rounded crystalline spar region (Fig. 8A).
Horizontal cross-sections through vertically accreting samples contain round molds and
round micrite remains of consistent size (~10 µm, ~50 µm, and ~100 µm in diameter; Fig. 9).
The porous regions also contain entombed fossils of ostracods and pennate diatoms (Fig. 10), a
feature previously noted by Pedone and Caceres (2002). Single diatoms are typically found
occupying pore spaces (Fig. 10A).
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Microstructure of banded carbonate
Microscopically, the banded carbonate is well laminated, consisting of laminae that are
either spar-dominated with minor micrite or micrite-dominated with minor spar. Fibrous spar
layers are laterally persistent and maintain a constant thickness (Figs. 11A–D), whereas micrite
layers vary in thickness along individual laminae and display microclots and filamentous
structures (see arrows in Figs. 11B–D). Laminations that are dominantly micritic are highly
clotted and uneven (Figs. 11E–F). Additionally, dense spar laminations can be traced
unambiguously from subunit to subunit, whereas the uneven nature of micrite laminations
complicates our ability to match them across subunits (Figs. 12–13).
Stable isotope results
Stable isotopic results of carbon and oxygen of the three fabric types are listed in Table 1.
The mean δ
13
C values for the banded micrite, banded spar, and porous fabric are −0.51 ± 0.21
(1σ, n = 9), −0.45 ± 0.26 (1σ, n = 9), and −0.50 ± 0.23 (1σ, n = 10), respectively. The mean δ
18
O
values of the banded micrite, banded spar, and porous fabric are −7.18 ± 0.41 (1σ, n = 9), −7.40
± 0.39 (1σ, n = 9), −7.53 ± 0.44 (1σ, n = 10), respectively. Figure 14 is a cross plot of the δ
18
O
and δ
13
C values from Table 1 and data from Becker et al., (2001), displaying a positive linear
trend with an R-value of 0.85. Figure 15 demonstrates the relative changes in δ
18
O and δ
13
C
clustered by fabric type.
DISCUSSION
Interpretation of Microfabrics
Porous Weakly Laminated Carbonate
Microfabrics of the porous regions display a range from microclots in a matrix of spar to
densely clotted micrite that makes up the primary framework (Figs. 5A–C). Dense micrite and
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clotted fabrics are common components of microbial carbonates, usually interpreted as the
remains of calcified cells or calcification associated with extracellular polymeric substances
(EPS) (Riding, 2000; Bontognali et al., 2008). Similar fabrics have been identified in other
lacustrine carbonates and are also interpreted as the products of microbial calcification (Arp,
1995; Pache et al., 2001; Nehza et al., 2009). Analyses of modern thrombolites from Kelly Lake,
British Columbia found coccoid cyanobacteria, small filamentous cyanobacteria, heterotrophic
bacteria, and diatoms to be the dominant contributors to the formation of clotted fabrics (Ferris et
al., 1997). Despite the lack of distinct morphological evidence for a microbial affinity, the
resemblance in texture for micrite that is associated with filamentous remains to areas that
contain calcified micritic shrubs and clots, provides further support for a microbial origin of the
microclots (Figs. 5E–D). Additionally, the occurrence of isopachous spar and microspar on
clusters of rounded micritic aggregates is a ubiquitous feature in freshwater tufas and travertines,
resulting from calcium carbonate precipitation on biofilms (Scholl and Taft, 1964; Pedley, 1987;
Freytet and Verrecchia, 2002). Microbially produced EPS creates diffusion limited sites that
affect adsorption and interactions with calcium ions, thereby influencing the mineral product
(Verrecchia, 1996).
Another common feature of the porous regions are pennate diatoms (Fig. 10A). Diatoms
are conspicuous components of lacustrine carbonates (Kempe et al., 1991; Ferris et al., 1997;
Arp et al., 1999b; Rosen et al., 2004), producing copious amounts of EPS, which are known to
favor both sediment trapping and mineral precipitation (Winsborough and Golubic, 1987;
Riding, 2000). The occurrence of diatoms within pore spaces suggests that EPS degradation
could have resulted in the development of the pores where aligned elongate fenestrea can result
due to the oxidation of seasonal diatom clumps (Monty, 1976). In some cases, diatoms are
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oriented so that their long axis is surface normal (Fig. 10A), indicating a potential phototropic
response (e.g., Cohn and Weitzell, 1996).
In all instances where filamentous microfossils were detected, multiple filaments were
found together forming distinct laminations (Fig. 7). Calcified filamentous microfossils with
similar orientation (perpendicular to subperpendicular relative to the laminations) occur in
several other tufa deposits and likely represent a phototropic behavior (Ferris et al., 1997; Arp et
al., 1999b; Valero-Garces et al., 2001; Arenas et al., 2007). Although photosynthesis-induced
carbonate precipitation could have resulted from CO
2
uptake through oxygenic photosynthesis, it
has been demonstrated that the effect of photosynthetic CO
2
fixation in promoting
supersaturation is not of primary importance in waters of high alkalinity (Arp et al., 1999a).
Instead, calcification associated with microbial EPS was likely a more influential factor on the
carbonate fabric. Specifically, porous tufa fabrics have been observed to result from
heterotrophic bacterial EPS degradation and associated gas formation (Arp et al., 1998).
Based on their size and orientation, the filaments possibly represent cyanobacteria of the
genus Oscillatoria, or Dichothrix, which are common in calcifying settings (Ferris et al., 1997;
Arp et al., 1999a; Planavsky et al., 2009), however, actual affinities are not discernible. Larger
round voids (~50–100 µm in diameter) are likely molds through bundles of algal filaments (Figs.
9C–D). In cases where molds do not occur, secondary micritic cements filled the voids.
Planavsky et al. (2009) and Turner et al. (2000) observed similar textural gradients in filament
preservation, where they suggest that initial precipitation took place on filament sheaths and
subsequent cement growth occurred in association with filament degradation.
Subspherical peloids (~100 µm in diameter) that exhibit isopachous spar cements are also
common in the porous regions. Peloids by definition are of unknown origin (McKee and
! 186
Gutschick, 1969) as determining their origin precisely is complicated by their inherent
clotted/opaque texture. At least three different possible origins exist for those described here: (1)
fecal pellets, (2) precipitation associated with bacterial colonies (Chafetz, 1986; Riding and
Tomás, 2006), and (3) abiotic origin (sensu Bosak et al., 2004).
Fibrous spar laminations of consistent lateral thickness are almost always associated with
micrite laminations of irregular thickness across a single lamination (Figs. 5F–G). Spar-micrite
couplets are widespread in freshwater fluvial tufas (Irion and Müller, 1968) and have also been
termed freshwater isopachous fringe cements (Pedley, 1987; Pedley, 1992). Using mesocosm
flume experiments that simulate fluvial tufa deposition, Pedley et al., (2009) demonstrate that
EPS dominated surfaces can form clotted peloidal spherulites upon which sparite crystals
develop. Their results suggest that EPS serves as a nucleation site upon which spar crystals
form, whereas UV treated controls did not produce the same results. While it is likely that the
dynamics that govern fluvial tufa formation might not be exactly analogous to those that form
lacustrine tufa (e.g., Arp et al., 2010), the widespread occurrence of spar-micrite couplets in
lacustrine and fluvial settings suggest a similar formation mechanism. The irregular thickness of
micrite layers and their association with clots suggest they could be the result of EPS/microbial
calcification (e.g., Ford and Pedley, 1996). Fringing (sometimes radial, Fig. 5G) spar layers
often succeed the micrite laminations, suggesting the micrite bands exert at least a passive
influence on the succeeding bands by serving as antecedent topography.
Dense Laminations
The spar-micrite layers that comprise the mesoscopic bands are similar to the isopachous
fringes from the weakly laminated carbonate, however, the spar in the banded regions is thicker,
! 187
extremely even in thickness, and more laterally continuous. The banding is so prominent that it
can easily be traced across the tufa subunits (Fig. 12).
The occurrence of the same dense laminations across tufa subunits presents compelling
evidence for a control at or beyond the meter scale mound (e.g., climate, spring activity, lake
chemistry) for the thick fibrous layers of the bands. On the other hand, the variability in
thickness and clotted nature of the thin micrite layers that separate the fibrous spar suggest a
dominance of microscale or “local” controls on their formation.
Here we suggest a microscale control on the formation of the dense micrite laminations
given their lateral variability in thickness at the microscale leading to an ultimate variability
across subunits (Figs. 11E–F; Fig. 13). In contrast to the fibrous spar laminations, the micritic
laminations are locally controlled at the mm scale, narrowing down their origin to submillimeter
factors (e.g., microbial influence, dissolution, local cement growth). These observations agree
with previous interpretations of uneven micrite laminations as lithification associated with the
presence of microbial mats (Monty, 1976; Riding, 2000).
Centimeter to millimeter bands that result from in situ precipitation are a common feature
of lacustrine carbonates (Pentecost, 2005; Arenas et al., 2007; Last et al., 2010). Light/dark
laminae often reflect seasonal variations in calcification and microbial growth (Ford and Pedley,
1996; Riding, 2000). However not all lacustrine carbonates are influenced by microbial activity
and demonstrating microbial influence in ancient deposits is complicated by poor preservation
and diagenesis. Given the frequent use of cm–mm scale banding to reconstruct ancient
environments, it is critical to attempt to narrow down the dominant controls on banding
formation and determine if they (1) indeed reflect primary signals and (2) are microbially
mediated, which can introduce vital effects (Andrews et al., 1993; Ferris et al., 1997; Sumner,
! 188
2001). Using spatial relationships with specific carbonate samples and identifying bands that
were deposited at the same time allows us to get an approximation of the dominant scales of
control on laminae formation.
Bands that match across subunits that originated from the same mounds (see Fig. 12)
suggest that the “events” that match are controlled at the scale of the tufa mounds or greater
(spring activity, climate, lake chemistry). On the other hand, the micrite laminations as well as
the porous weakly laminated carbonate demonstrate significant variability in texture across
samples, reflecting a dominance of local influence at the cm-mm scale level by microscale
processes (e.g., microbial processes, local cement growth). Furthermore, successfully
identifying samples that formed far away from one another (e.g., opposite sides of a lake) but
along strike and that record the same “events” would imply that the portions that match represent
nonlocal controls on formation—controlled at the scale of the lake itself (e.g., climate, lake
chemistry).
Stable isotope interpretations
Observations on the stable isotopic values of the various fabrics investigated indicate
there is little variability across the different fabric types (Table 1, Fig. 15). These results are
consistent with other work that demonstrates that in environments with a large dissolved
inorganic carbon (DIC) pool, typical of alkaline settings, the effect of photosynthetic CO
2
removal on δ
13
C isotopic values is negligible (Arp et al., 2001). Instead it likely that biofilms
contributed to altering the fabric by playing a passive role via by diffusion-controlled and EPS-
mediated mineralization (e.g., Bontognali et al., 2008). Thus, although petrographic
observations reveal morphological biological influence (Fig. 5), there is a lack of a chemical
signal of a potential metabolic effect on the carbonates. The co-variation of δ
13
C and δ
18
O
! 189
values shown in Figure 14 further demonstrates that lake Barstow was a closed basin lake given
the high R-value characteristic of closed lacustrine systems (Talbot, 1990).
Multiscale Controls on Tufa Formation
Considered collectively, the above observations and interpretations reflect a dynamic
interplay of both environmental and microbial controls on tufa formation. At a macroscale, tufa
mounds of the Barstow Formation formed when Ca-rich groundwater entered the saline alkaline
lake via spring orifices. The bicarbonate rich water would have fostered rapid localized
carbonate formation at sites where groundwater entered the lake. The presence of vertical tubes
in the unit directly underlying the tufas lends support for a localization mechanism provided by
spring activity (Becker et al. 2001). Localization of carbonate mounds around springs is a
common tufa formation mechanism observed in many modern and ancient lakes (Scholl, 1960;
Kempe et al., 1991; Thompson et al., 1990; Bischoff et al., 1993; Arp et al., 1999a; Rosen et al.,
2004).
At a mesoscale, the occurrence of tubular voids (Fig. 3C) and tubular secondarily filled in
molds (Fig. 3D) suggest that aquatic plants served as suitable sites for nucleation as reported by
Pedone and Caceres (2002). Nucleation around aquatic plants is a common feature in lacustrine
carbonate deposits (Riding, 1979; Arenas et al., 2000; Arenas et al., 2007). However, given that
not all of our samples display tubular molds (see Fig. 3E) indicates that while suitable for
nucleation, twigs and roots of aquatic plants were not a requirement for carbonate nucleation and
tufa formation.
The link between mesoscopic macrophyte encrustations and a component of microscale
controls is elucidated via a close up view of the tubular molds. In areas where carbonate
nucleated onto putative twigs or roots of aquatic plants, a microbial fabric surrounds the void
! 190
spaces as if filaments once radiated from the nucleation surface (Figs. 7C and 8A). The
microbial dominance in framework between the mold and the rest of the tufa subunit indicates
that microbial biofilms possibly facilitated the nucleation of subsequent carbonate precipitation
(Fig. 7C).
Microscopically, the porous weakly laminated carbonate is dominated by clotted peloidal
micrite in matrix of microspar or spar. The occurrence of filamentous microfossils, round voids
of consistent size in horizontal cross section, micritic shrubs and clots that likely represent
microbial remains, and the abundance of diatoms and ostracods, present strong evidence for a
significant microbial component in the development of the porous mesofabrics. Additionally,
laminoid fenestrae associated with micrite laminations are similar to irregular fenestral fabrics
common in microbial carbonates that could be the result of detachment of surficial algal mat
layers from underlying ones, microbial matter oxidation/decay, and/or fenestra associated with
oxygenic photosynthesis that leaves behind elongate, irregular, horizontal or vertical cavities
(Golubic, 1973; Monty, 1976; Arp et al., 1998; Mata et al., 2012). The weakly laminated regions
likely reflect couplets of microbial micrite precipitation and abiotic spar formation.
Mesoscopic laterally continuous dense bands of constant thickness (Figs. 11A–B)
represent largely abiotic spar growth that reflect carbonate precipitation controlled at the scale of
the m-scale tufa mound or greater (spring activity, lake chemistry, climate). On the other hand,
uneven micrite laminations are more strongly controlled by local cm–mm mechanisms (e.g.,
microbial activity, local cement growth).
These observations indicate that the size and dominant features of the product do not
necessarily scale with the scope of the process. For instance, the remarkably consistent
representation of the same dense laminations across multiple subunits indicates that meter-scale
! 191
processes or greater dominate local processes (e.g., microscale controls). Conversely, microbial
features comprise a significant portion of the clotted/porous regions, in some cases forming the
primary framework and possibly inducing mesoscopic pores, yet are largely being governed at
the micro-mm scale. Similarly, the morphology and shape of several of the tufa subunits is
dictated by macrophyte encrustations, propagating the initial shape long after the plant has been
encrusted. This disparity in scale of the product versus the span and extent of control of the
process further highlights the importance and value of (1) analyzing multiple subunits, and (2)
conducting multiscale analyses (e.g., Shapiro, 2000). Failure to recognize this potential
divergence in scale of control versus scale of the feature might lead to inaccurate interpretations
of the textures and their corresponding environmental signals.
Implications for Early Earth Microbialites
Analyses described here indicate that the Barstow Formation tufa formed via in situ
precipitation of carbonate as indicated by the presence of calcified fossils in life position as well
as the dominance of isopachous (fringing) laminae. Growth via in situ precipitation of minerals
was widespread in the early Earth (Grotzinger and Knoll, 1999; Riding, 2011). Modern marine
microbialites, on the other hand, result largely from the trapping and binding of grains resulting
in coarse agglutinated deposits that rarely exhibit fine laminae (Reid et al., 2000; Dupraz and
Visscher, 2005). Therefore, although microbialites occur in some modern marine environments,
they do not appear to be forming via the same mechanisms that resulted in the formation of their
ancient counterparts (Awramik and Riding, 1988; Fairchild, 1991; Arp et al., 1999a; Grotzinger
and Knoll, 1999; Riding, 2008). Conversely, lacustrine carbonates closely resemble
Precambrian microbialites in several aspects: (1) formation via in situ mineralization as opposed
to trapping and binding of loose sediment (2) their ability to form the entire range of mesofabrics
! 192
common in ancient deposits (e.g., dendritic, thrombolitic, stromatolitic) and (3) their propensity
to form large decameter domes and towers (Scholl, 1960; Riding, 1979; Kempe et al., 1991), a
characteristic common in the Precambrian but not as widespread in modern marine settings.
Furthermore, laminated lacustrine carbonates like those described in this contribution, are
strikingly similar to Precambrian stromatolite crust fabrics that range from spar crusts (Fig.
11A), hybrid crusts (Fig. 11B), and micrite crusts (Figs. 11E–F), respectively reflecting a
dominantly abiotic to biotic gradient of control in fabric formation (Riding, 2008).
Therefore, although they did not form in the same environment, lacustrine carbonates are
suitable textural and morphological analogs that might aid our interpretation of ancient deposits
with similar textures. Understanding the relative contribution of biotic and abiotic components
and being able to thoroughly assess how these factors manifest into meso and macrosfabrics will
allow us to (1) more accurately exploit/understand their environmental signatures, (2) ascertain
biosignature presence and preservation and (3) improve our methods for assessing biogenicity at
a meso and macroscale given the potential for discovering similar deposits elsewhere in the Solar
System.
CONCLUSIONS
We investigated lacustrine tufa from the Middle Miocene Barstow Formation to assess
their mechanisms of formation and determine how various formation mechanisms translate into
their texture and morphology. Tufa deposits resulted from multi-process and multi-scale levels
of control that each left behind a unique trace on their morphology. Macroscopically, the m-
scale tufa mounds formed when groundwater fed the saline alkaline lake via localized spring
orifices imparting a localized development of the carbonate mounds. Some tufa subunits
nucleated onto stems or roots of aquatic plants resulting in imprints of tubular molds or
! 193
secondarily filled voids. The filamentous nature of the clotted fabric that immediately surrounds
the molds indicates that microbial biofilms assisted in the nucleation and growth of the tufa
subunit. The weakly laminated porous carbonate is dominated by a clotted/peloidal fabric and
also contains abundant diatoms, ostracods, filamentous remains, and fenestral laminations—all
indicative of a significant microbial component.
Dense spar laminations are controlled by nonlocal processes or at least at the scale of the
mound itself (e.g., spring activity, climate, lake chemistry). The variability in texture across
subunits for micritic laminations indicates their fabric is dominantly locally controlled at the cm-
mm scale (e.g., microbial activity, local cement growth). The tufa subunits resulted from in situ
precipitation of oversaturated waters with respect to calcium carbonate, revealing textures that
further demonstrate their striking textural resemblance to ancient microbialites.
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! 203
FIGURES
Figure 1. Geologic map of study site. The stars denote the three tufa-bearing sites. Modified
from Cole et al. 2005.
! 204
Figure 2. Outcrop images of sampled tufa mound at the Owl Campground tufa-bearing site. (A)
Image of tufa mound, hammer for scale. (B) In situ tufa subunits displaying concentric
laminations.
! 205
Figure 3. Mesostructure of tufa subunits. (A) Stromatolitic tufa. (B) Nodular tufa. (C-D)
vertical cross sections of tufa subunits, note the tubular internal molds. (E) Vertical cross section
of a stromatolitic tufa subunit.
! 206
Figure 4. Mesoscopic textures. (A) Polished slab of a tufa subunit with the three main textures
labeled. (B) Corresponding thin section photomicrograph of (A).
! 207
Figure 5. Microstructure of the weakly laminated, porous regions. (A-C) Micritic, peloidal
aggregates of varying density in a matrix of spar and microspar. (D) Micritic shrubs. (E)
Filamentous, micritic molds. (F-G) Couplets of irregular discontinuous laminae of spar and
micrite. Blue and white areas denote pore space.
! 208
Figure 6. Laminoid fenestrae. (A-B) Laminoid fenestrae of the porous, weakly laminated
regions.
! 209
Figure 7. Oriented filamentous microfossils. (A-B) Micritic filamentous molds in a matrix of
microspar from stromatolitic tufa subunit. (C-E) Micritic filamentous microfossils oriented
surface normal in a matrix of micrite.
! 210
Figure 8. Dense filamentous micrite fabric from the porous, weakly laminated regions. (A)
Dense filamentous micrite radiating from a secondarily filled mold in cross section. (B)
Filamentous micritic framework.
! 211
Figure 9. Cross sections of round molds. (A) Cross sections of filamentous micritic molds ~10
µm in diameter, inset denotes a mold (blue pore space) next to a structure filled with micrite. (B)
Cross sections of filamentous molds ~50 µm in diameter. (C-D) Round micritic structures of
consistent size, ~ 100 µm in diameter.
! 212
Figure 10. Microfossils of the porous regions. (A) Oriented diatoms, in dashed circles. (B)
Pennate diatom. (C) Ostracod fossil. Blue and white areas denote pore space.
! 213
Figure 11. Photomicrographs of banded carbonate. (A) Banded, fibrous calcite. (B) Banded
fibrous calcite, with minor micrite, arrows denote uneven thickness of micrite layer. (C)
Transition from banded spar to porous, weakly laminated region, arrows denote microclots in
micrite layers. (D) Micritic band with putative filamentous microfossils. (E-F) Micrite bands
with minor spar, note the uneven thickness of the micrite bands, arrow in (F).
! 214
Figure 12. Mesostructure of four subunits with the bands aligned. Scale bar = 2 mm.
! 215
Figure 13. Photomicrographs of three tufa subunits with the bands aligned. Note the continuous
nature of the spar laminae vs. the irregular nature of the micritic laminae.
! 216
Figure 14. Carbonate δ
13
C and δ
18
O cross plot.
18
O
13
C
data from Becker et al., 2001
porous carbonate
banded spar
banded micrite
y = 4.543 + 0.674x R = 0.8533
! 217
Figure 15. Carbonate stable isotopes of carbon and oxygen by fabric type.
18
O 13
C
banded spar
banded micrite
porous, weakly
laminated carbonate
banded spar
banded micrite
porous, weakly laminated carbonate
! 218
Table 1. Stable isotopic compositions of δ
18
O and δ
13
C (‰ VPDB) for carbonate from their
corresponding fabric type.
Fabric Type Sample δ
13
C δ
18
O
Micrite Bands
BF-RMR -0.528 -7.042
BF-RML -0.357 -6.964
BF-RMM -0.639 -7.237
BF-3Mb -0.578 -7.334
BF-1Mb -0.765 -7.504
BF-1Mb1 -0.608 -7.536
BF-1-ML -0.589 -7.477
BF-R-MR -0.020 -6.235
BF-R-ML -0.513 -7.267
Micrite Mean -0.511 -7.177
1σ 0.214 0.405
Spar Bands
BF-1-SR -0.521 -7.087
BF-1-SM -0.493 -7.465
BF-3-SL -0.706 -7.841
BF-1-SL -0.427 -7.075
BF-R-SL -0.339 -7.465
BF-R-SM -0.625 -7.882
BF-3Sb -0.240 -7.290
BF-RSL 0.081 -6.760
BF-RSR -0.772 -7.772
Spar Mean -0.449 -7.404
1σ 0.261 0.387
Porous Fabric
BF-1C1 -0.428 -7.145
BF-3C1 -0.772 -7.574
BF-3C2 -0.331 -7.260
BF-3C3 -0.549 -7.688
BF-3C4 -0.297 -6.988
BF-FC1 -0.568 -8.255
BF-FC2 -0.837 -8.251
BF-FC3 -0.088 -7.131
BF-FC4 -0.643 -7.486
BF-RC1 -0.445 -7.482
Porous Fabric Mean -0.496 -7.526
1σ 0.227 0.441
Mean Overall -0.486 -7.374
1σ 0.227 0.438
Abstract (if available)
Abstract
Microbialites, defined as organosedimentary deposits that accreted as a result of benthic microbial activity, are widely recognized as potential records of ancient microbial life and environmental archives. However, deciphering the relative role of the various mechanisms that result in their morphogenesis remains a challenge, yet is critical if we are to fully unravel their potential as biosignatures and environmental archives throughout the geological record. In this dissertation, three case studies focusing on the environmental effects on microbialite morphogenesis are presented. In particular, the spatial, temporal, and depositional settings are investigated to understand how the scales of control (local, regional, global) affect microbialite growth, and in turn, affect microbialite veracity as environmental recorders. ❧ Spring-fed fluvial carbonate environments can host microbialites when carbonate springs (cold or hot) emerge at the surface and form carbonate deposits that are typically heavily encrusted in microbial remains. Fluvial carbonates that are sourced from cold water may be important indicators of past pluvial periods and thus can be useful in climate reconstructions. Currently, few fluvial tufa records exist from southern California—a region projected to experience significant aridification within the coming decades (Seager et al., 2007)—however, to date, fluvial tufa records have not been described in coastal southern California, close to most modern metropolitan centers. Recently discovered spring-associated carbonates from coastal southern California, located near Zaca Lake, were investigated. Petrographic study revealed the presence of the calcite biosignature of the desmid microalgae Oocardium stratum, a microalgae known to occur exclusively in carbonate-depositing systems sourced from ambient-temperature water (~9-13 °C) and thus confirming the carbonates potential as indicators of past pluvial. Radiocarbon and Infra-Red Stimulated Luminescence (IRSL) of the carbonates suggest at least two episodes of carbonate growth: one at 19.4 ± 2.4 through 17.8 ± 2.8 and another at 11.9 ± 1.5 ka verified with a charcoal ¹⁴C age of 10.95 ± 0.12 cal ka BP. The wet period indicated by tufa growth between 19.4 and 17.8 is relatively consistent with other southern California climate records both north and south of Zaca Lake. Tufa growth ca. 12 to 11 ka demonstrates wet conditions occurred as far south as Zaca Lake, which is in contrast to climate records south of this site in Lake Elsinore that indicate persistently dry conditions through this interval. This work suggests that rather small shifts in the average winter season storm track could produce large changes in regional hydroclimate. ❧ Remarkably aerially extensive Upper Triassic carbonate microbialites from the southwestern United Kingdom (SW UK) occur within the Triassic-Jurassic boundary interval, but have always been considered decoupled from the end-Triassic mass extinction event. Their vast lateral extent makes them amenable to study the influence of scale of process (local, regional, global) on microbialite morphogenesis (discussed below) as well as their significance to the Triassic-Jurassic mass extinction. The microbialites (1) occur at the same stratigraphic level as the mass extinction (2) contain a negative carbon isotope excursion in δ¹³Corg and (3) co-occur with a bloom of ‘disaster taxa’ prasinophyte algae also dominant in other European sections in synchrony with a widely observed carbon isotope excursion. Further, the microbialites are composed of intricate microbial mat textures and correlate with mixed carbonate-siliciclastic facies that contain abundant, well-preserved filamentous microfossils whose unusual preservation implies growth in waters with anomalously high carbonate saturation. High saturation state with respect to calcium carbonate may have resulted from a period of enhanced continental weathering in response to higher pCO₂ levels triggered by the emplacement of the Central Atlantic Magmatic Province (CAMP). These findings indicate the end-Triassic strata of the SW UK capture a significant microbial carbonate sedimentary response as is commonly observed across other episodes of biotic crisis. ❧ A major challenge in the study of microbialites throughout the rock record is disentangling the complex mix of processes (local and nonlocal) that influence their morphology. By assessing the degree to which microbialite textures vary laterally along strike, however, it may be possible to constrain the scales of the processes of control that impose the greatest influence on their morphology. In one example, analyses of microbialite tufa (carbonate) mounds from the lacustrine sediments of the Miocene Barstow Formation reveal multi-scale and multi-process controls on their formation that each left behind a unique imprint on their morphology: m-scale localization caused by spring activity, cm-scale sites for nucleation provided by aquatic plants, and sub-mm-scale changes in texture. The presence of microfossils and a ubiquitous clotted micritic texture that varies in thickness from sub-unit to sub-unit, suggests microbial influence was ‘locally’ controlled at the cm-scale. In contrast, the spar laminae maintain a constant thickness across sub-units, suggesting control at or beyond the m-scale mound likely representing larger scale changes in lake chemistry, climate, and/or spring activity. Combined multi-scale (sensu Shapiro, 2000) and lateral continuity observations from the three case studies in this dissertation reveal that: (1) lateral patterns of textural continuity are strongly dependent on the nature of the depositional system, (2) some features are controlled by local processes whereas others are primarily controlled by regional processes, (3) in some instances, the size/scale of the feature does not necessarily scale with the scope of the process—important for addressing at which scales measurements should be done and enabling us to predict the relative contribution of regional and local changes in the rock record. Although this method does not reveal explicit processes of control it can be used to determine the relative scale of the significant process(es) of control and in turn inform our understanding of the environmental scope of the textures under investigation.
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Ibarra, Yadira
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The geobiology of fluvial, lacustrine, and marginal marine carbonate microbialites (Pleistocene, Miocene, and Late Triassic) and their environmental significance
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